Large Ecosystem Perturbations: Causes and Consequences (GSA Special Paper 424)

Large Ecosystem Perturbations: Causes and Consequences

edited by Simonetta Monechi, Rodolfo Coccioni, and Michael R. Rampino

THE GEOLOGICAL SOCIETY OF AMERICA® Special Paper 424

Copyright © 2007, The Geological Society of America, Inc. (GSA). All rights reserved. GSA grants permission to individual scientists to make unlimited photocopies of one or more items from this volume for noncommercial purposes advancing science or education, including classroom use. For permission to make photocopies of any item in this volume for other noncommercial, nonprofi t purposes, contact the Geological Society of America. Written permission is required from GSA for all other forms of capture or reproduction of any item in the volume including, but not limited to, all types of electronic or digital scanning or other digital or manual transformation of articles or any portion thereof, such as abstracts, into computer-readable and/or transmittable form fo r personal or corporate use, either noncommercia l or commercial, for-profit or otherwise. Send perm ission requests to GSA Copyright Permissions, 3300 Penrose Place, P.O. Box 9 140, Boulder, Colorado 80301-9 140, USA. Published by The Geological Society of America, Inc. 3300 Penrose Place, P.O. Box 9 140, Boulder, Colorado 8030 1-9 140, USA www.geosociety.org Printed in U.S.A. GSA Books Science Editors: Marion E. Bickford and Abhijit Basu Library of Congress Cataloging-in-Publication Data Large ecosystem perturbations : causes and consequences I edited by Simonetta Monechi, Rodo lfo Coccioni, Michael R. Rampino. p. em. - (Special paper; 424) Includes bibliographical references. ISBN 978-0-8 137-2424-9 (pbk.) I . Paleoecology- Congresses. 2. Paleoclimatology-Congresses. 3. Global environmental change- Congresses. 4. Extinction (B io logy)-Congresses. I. Monechi, Simonetta. ll. Coccioni, R. Ill. Rampino, Michael R. QE720.L37 2007 560'.45-dc22 20070 12987 Cover, above: Panoramic view of the Zumaia section (Basque Coun try, Spain), one of the most expanded deep-water successions across the upper Paleocene and lower Eocene transition. Photograph taken by Eustoquio Molina. Below left: Cross-polarized light micrograph of a smear sl ide showing a calcareous nannofossil assem blage mai nl y consisting of Toweius and Zygrhablithus at the PaleoceneEocene transition, Site 690. Photograph taken by Simonetta Monechi. Below middle: Cross-polarized light micrograph of a smear slide showing a calcareous nannofossil assem blage mainl y consisting of Discoaster and Fasciculithus at the Paleocene-Eocene transition, Site 690. Photograph taken by Simonetta Monechi. Below right: Cross-polarized light micrograph of a smear slide showing a calcareous nannofossil assem blage mai nly consisting of Zygrhablithus, Chiasmolithus, Fasciculithus, and Toweius at the Paleocene-Eocene transition, Site 690. Photograph taken by Simonetta Monechi .

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Contents Preface . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . v 1. Cenozoic mass extinctions in the deep sea: What perturbs the largest habitat on Earth? . . . . . . . . . 1 Ellen Thomas 2. A major Pliocene coccolithophore turnover: Change in morphological strategy in the photic zone . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 25 Marie-Pierre Aubry 3. The Paleocene-Eocene Thermal Maximum in Egypt and Jordan: An overview of the planktic foraminiferal record . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 53 Elisa Guasti and Robert P. Speijer 4. Calcareous nannofossil assemblages and their response to the Paleocene-Eocene Thermal Maximum event at different latitudes: ODP Site 690 and Tethyan sections . . . . . . . . . . . 69 Eugenia Angori, Gilen Bernaola, and Simonetta Monechi 5. A review of calcareous nannofossil changes during the early Aptian Oceanic Anoxic Event 1a and the Paleocene-Eocene Thermal Maximum: The influence of fertility, temperature, and pCO2 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 87 Fabrizio Tremolada, Elisabetta Erba, and Timothy J. Bralower 6. Ecosystem perturbation caused by a small Late Cretaceous marine impact, Gulf Coastal Plain, USA . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 97 David T. King Jr., Lucille W. Petruny, and Thornton L. Neathery 7. Chemostratigraphy of Frasnian-Famennian transition: Possibility of methane hydrate dissociation leading to mass extinction . . . . . . . . . . . . . . . . . . . . . 109 Mohammad Hossein Mahmudy Gharaie, Ryo Matsumoto, Grzegorz Racki, and Yoshitaka Kakuwa

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Preface

The history of life on our planet has been punctuated by many dramatic and widespread biological variations. Mass extinctions, rapid turnovers and evolutionary novelties are often associated with large-scale climate changes, geological events and extraterrestrial impacts which have influenced and accelerated events and processes of the biological systems. The interplay between major environmental perturbations and faunal and floral modifications is best seen in the marine record which provides a very detailed, often complete, documentation of paleobiological changes. This volume contains a series of papers on the marine record which were presented during the International Geological Congress of Florence in August 2004 in the topical session T04 “Mass-extinctions and large ecosystem perturbations: Terrestrial versus extraterrestrial causes.” The volume also includes papers presented at the workshop WSB01 “Past and future contributions of nannoplankton research to global change.” Contributions deal with different groups, various environments, and different time intervals. The paper presented by Thomas shows how deep-sea benthic foraminifera underwent severe extinction and faunal turnovers during the Cenozoic but, surprisingly, did not suffer significant extinction at the K/Pg, when the plankton communities experienced catastrophic destructions. Several papers focus on the Paleocene-Eocene transition, which was characterized by major turnovers in the phytoplankton and planktonic foraminiferal communities and important extinctions in the deep-sea environment associated with the Paleocene-Eocene Thermal Maximum (PETM). Guasti and Speijer have studied the turnover in planktonic foraminiferal assemblages across the PETM in the Western Tethys, stressing the taphonomic effect (dissolution) as a cause for the major bloom of Acarinina which is usually argued to be indicative of nutrient influxes in a perturbed marine environment. The paper by Angori, Bernaola, and Monechi reports on the calcareous nannofossil record from different environments and latitudes showing two main sets of regimes: increases of r-selected taxa, probably in response to upwelling and nutrient influx before the PETM and the inception of the Carbon Isotope Excursion (CIE), followed by a rapid expansion of warm-water assemblages into higher latitudes (e.g., ODP Site 690) induced by a rise of surface-water temperatures. Tremolada, Erba, and Bralower compare changes occurring in calcareous nannofossil during the early Aptian Oceanic Anoxic Event 1a with those of the PETM. Both were characterized by extreme greenhouse regimes. Yet, the nannofloral responses were opposite. Opportunistic eutrophic taxa thrived in the Aptian, with the concomitant demise of the oligotrophic nannoconids, whereas the PETM conditions (global nutrientdepleted surface waters) favoured oligotrophic-mesotrophic taxa. Other topics discussed concern various aspects of ecosystem variations. Aubry delineates the hypothesis of an overall evolutionary strategy characterized by a morphological shift of the Coccolithophores toward smaller size. This trend, already under way in the Neogene, becomes fully manifest at the end of the Pliocene, concomitant with the inception of the Pleistocene global cooling. Extraterrestrial or internally driven perturbations inducing environmental changes affect the biota represented at different scales. The contribution of King, Petruny, and Neathery describes the aftermath of a minor extraterrestrial impact in the Gulf Coastal Plain, USA, at 83.5 m.y. and its relatively minor paleobiological effects, as expected without global or even regional extinctions. Mahmudy Gharaie, Matsumoto, Racki, and Kakuwa, propose a possible mechanism for the severe mass extinction which took place in the Late Devonian. On the basis of stable isotopes data from China and the Urals, rapid increase of ocean temperatures is assumed. The effects of this warming were enhanced v

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Preface by flood-basalts, gaseous effusions, and sea-level drop, with an abrupt release of vast marine methane hydrates into the hydrosphere and atmosphere. The papers contained in this volume offer case studies on a large spectrum of ecosystem changes with various causes and consequences, and provide stimulating ideas for future works. Many thanks to the reviewers for their perceptive critique and helpful advice. Reviewers who agreed to be identified are: Laia Alegret, Laurel Bybell, Luc Beaufort, Simone Galeotti, Andy Gooday, Lee Kump, Maurits Lindstrom, Eustoquio Molina, Joerg Mutterlose, David Powars, Murice Tucker, Tyler Volk, Paul Wignall, and Woody Wise. Thanks also for the assistance of Guia Morelli, Flavia Tori, and Alessandro Mari Catani. We acknowledge the support of the Dipartimento di Scienze della Terra, University of Florence, Italy, and of grant MIUR/PRIN COFIN 2005 to S. Monechi. Abhijit Basu, Pat Bickford, Sarah Barkin, and the editorial staff of the Geological Society of America kindly provided the editorial fine-tuning necessary for bringing this volume to publication. To these individuals and organizations we offer many sincere thanks. Simonetta Monechi

The Geological Society of America Special Paper 424 2007

Cenozoic mass extinctions in the deep sea: What perturbs the largest habitat on Earth? Ellen Thomas* Center for the Study of Global Change, Department of Geology and Geophysics, Yale University, New Haven, Connecticut 06520-8109, USA

ABSTRACT Deep-sea benthic foraminifera live in the largest habitat on Earth, constitute an important part of its benthic biomass, and form diverse assemblages with common cosmopolitan species. Modern deep-sea benthic foraminiferal assemblages are strongly influenced by events affecting their main food source, phytoplankton (a relationship known as bentho-pelagic coupling). Surprisingly, benthic foraminifera did not suffer significant extinction at the end of the Cretaceous, when phytoplankton communities underwent severe extinction. Possibly, bentho-pelagic coupling was less strong than today in the warm oceans of the Cretaceous–Paleogene, because of differences in the process of food transfer from surface to bottom, or because more food was produced below the photic zone by litho-autotrophs. Alternatively, after the end-Cretaceous extinction the food supply from the photic zone recovered in less time than previously thought. In contrast, deep-sea benthic foraminifera did undergo severe extinction (30%–50% of species) at the end of the Paleocene, when planktic organisms show rapid evolutionary turnover, but no major extinction. Causes of this benthic extinction are not clear: net extinction rates were similar globally, but there is no independent evidence for global anoxia or dysoxia, nor of globally consistent increase or decrease in productivity or carbonate dissolution. The extinction might be linked to a global feature of the end-Paleocene environmental change, i.e., rapid global warming. Cenozoic deep-sea benthic faunas show gradual faunal turnover during periods of pronounced cooling and increase in polar ice volume: the late Eocene–early Oligocene, the middle Miocene, and the middle Pleistocene. During the latter turnover, taxa that decreased in abundance during the earlier two turnovers became extinct, possibly because of increased oxygenation of the oceans, or because of increased seasonality in food delivery. The Eocene-Oligocene was the most extensive of these turnovers, and bentho-pelagic coupling may have become established at that time. Keywords: deep-sea benthic foraminifera, extinction, K/Pg boundary, P/E boundary, E/O boundary, global warming, global cooling.

*Also at: Department of Earth and Environmental Sciences, Wesleyan University, Middletown, Connecticut 06459-0139, USA. Thomas, E., 2007, Cenozoic mass extinctions in the deep sea: What perturbs the largest habitat on Earth?, in Monechi, S., Coccioni, R., and Rampino, M.R., eds., Large Ecosystem Perturbations: Causes and Consequences: Geological Society of America Special Paper 424, p. 1–23, doi: 10.1130/2007.2424(01). For permission to copy, contact [email protected]. ©2007 The Geological Society of America. All rights reserved.

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INTRODUCTION The deep-sea floor (bathyal and abyssal depths, i.e., depths >200 m) constitutes the largest habitat on Earth (e.g., Norse, 1994; Verity et al., 2002). The deep-ocean floor had long been thought to be devoid of life (Wyville Thomson, 1873), but in the early 1840s the Antarctic expedition of James Clarke Ross on the Erebus found indications that life existed there. Convincing evidence emerged during the Lightning and Porcupine expeditions in the late 1860s (Wyville Thomson, 1873), which led to the organization of the Challenger Expedition (1872–1876; Murray, 1895). This expedition abundantly documented that the deep-ocean floor was inhabited by many forms of life, and one of the zoological volumes in the expedition reports describes deepsea foraminifera (Brady, 1884). Until the 1960s, deep-sea faunas were considered to have relatively low diversities, but studies in the second half of the twentieth century established that deepsea ecosystems are characterized by low density of individuals at high species diversities (e.g., Sanders et al., 1965; Grassle and Maciolek, 1992; Gage, 1996, 1997; Rex et al., 1997; Levin et al., 2001; Snelgrove and Smith, 2002). The deep sea is still one of the least known ecosystems on Earth, however, and even general features of its high diversity, comparable to that in tropical rain forests and coral reefs, are not well known or understood, e.g., whether there are diversity gradients between high and low latitudes (e.g., Gage, 1996; Culver and Buzas, 2000; Rex et al., 2000; Levin et al., 2001), and how to explain diversity gradients with depth (see discussion in Rex et al., 2005). In addition, we do not know how much of the organic matter in the oceanic carbon cycle is contributed by photosynthesizing eukaryotes, how much by photosynthesizing prokaryotes (e.g., Kolber et al., 2000), or how much by lithoautotrophic prokaryotes (e.g., Herndl et al., 2005; Bach et al., 2006). Present deep-ocean biota live perpetually in the dark, over most of the present ocean floor at temperatures close to freezing, under high pressures, at constant salinities, and in a world where very little food arrives, mainly derived from surface primary productivity hundreds to thousand of meters higher in the water column (e.g., Tappan, 1986; Gooday, 2003). Deep-sea biota outside hydrothermal vent regions and cold seep areas are thus surviving in an extremely low food environment, and are generally slow-growing and small. The seemingly monotonous environment is patchy on varying time-space scales, with skeletons of agglutinated, tree-shaped unicellular organisms (Xenophyophorans) serving as substrate for smaller unicellular biota (Hughes and Gooday, 2004), and with arrival of phytodetritus patches (algal debris aggregated by mucus from various organisms) in regions of seasonal blooms (e.g., Rice et al., 1994), as well as such unpredictable events as whale-falls (e.g., Gooday and Rathburn, 1999; Gooday, 2002; Smith and Baco, 2003; Rathburn et al., 2005). Environmental heterogeneity is enhanced because the lack of physical disturbance of the environment allows small morphological features of the seafloor to persist for long periods (e.g., Gage, 1996).

In this oligotrophic world small organisms far outnumber larger ones (e.g., Gage and Tyler, 1991; Gage, 1996). Abundant among the small life forms are the eukaryotic, unicellular foraminifera, many of which form a shell or test from organic matter, secreted calcium carbonate, or agglutinated sedimentary particles. Foraminifera are one of the most ecologically important groups of marine heterotrophic protists. Their history goes back to the Early Cambrian (Pawlowski et al., 2003), and they occur throughout the oceans, even in the deepest trenches (Todo et al., 2005). Species abundant in the fossil record are characterized by robust calcium carbonate or agglutinated tests, and are dominantly between 0.1 and 1 mm in size, with most forms in the meiofauna (0.1–0.3 mm). Some delicate agglutinated taxa, the komokiaceans, form a branching, tubular system, typically 1–5 mm across (Gooday et al., 1997), and are classified among the macrofauna of the deep sea. At depths greater than ~1000 m, foraminifera constitute more than 50% of the total eukaryotic biomass, with estimates of >90% at depths >2000 m in some regions (Gooday et al., 1992, 1998). Their assemblages are highly diverse: ~9000–10,000 living species have been described, with species determinations based on test morphology and composition (e.g., Goldstein, 1999). The shelled species are much better known than the naked (Pawlowksi et al., 1999) or organic-walled (soft-shelled) species, most of which have not yet received formal species names (e.g., Cedhagen et al., 2002). Studies of genetic material (e.g., Pawlowski and Holzmann, 2002) suggest that biological species may be more numerous than morphological species, with—as in other marine invertebrates—the few studied groups of foraminifera containing cryptospecies (e.g., Gage, 1996). Many of the morphological species are cosmopolitan in the present-day oceans. Benthic foraminifera move at speeds of only micrometers per hour (e.g., Kitazato, 1988; Gross, 1998; 2000), and almost certainly could not move at the speed suggested by the timing of first appearance of some species (e.g., Cibicidoides wuellerstorfi) in various oceans (Thomas and Vincent, 1987, 1988). Propagules formed during reproduction may be easily transported by ocean currents (Alve, 1999; Alve and Goldstein, 2003) and assist in the dispersal of species as well as in maintaining gene flow between distant populations. In such a large habitat, physico-chemical parameters do not change rapidly over the whole habitat, and large-scale, isolating barriers are absent (Gage, 1996). The easily transported propagules ensure that new populations will become established at locations distant from existing ones as soon as environmental conditions are favorable, so that disturbed regions become recolonized quickly (Kuhnt, 1992; Hess and Kuhnt, 1996; Hess et al., 2001; Alve and Goldstein, 2003). Under such circumstances, one would expect morphological species to have relatively long species lives, as indeed observed in deep-sea benthic foraminifera, which have average species lives of ~15 m.y. (Culver, 1993). Some common Recent morphological species, e.g., Oridorsalis umbonatus, have persisted since at least the Late Cretaceous (Kaiho, 1998).

Cenozoic mass extinctions in the deep sea Thus a major question is what caused mass extinctions in the deep oceans: What type of environmental disturbance could be so all-encompassing, major, and rapid that cosmopolitan deepsea species would suffer mass extinction, without the possibility to migrate vertically or horizontally and to repopulate from refugia? In this paper I use deep-sea benthic foraminifera, which of all deep-sea organisms have the most abundant fossil record, to probe this question; emphasis is on assemblages dominated by forms with secreted CaCO3 tests and agglutinated forms with CaCO3 cement, which in the present oceans occur at depths between several hundreds and ~4000–4500 m with the exception of the polar oceans (Gooday, 2003). I review our knowledge of present deep-sea benthic foraminiferal ecology, give an overview of deep-sea faunal assemblages through the Cenozoic, and discuss the decoupling between mass extinctions in the pelagic and benthic realms at the Cretaceous-Paleogene and the Paleocene-Eocene boundaries (Thomas, 1990b; Kaiho, 1994a). RECENT DEEP-SEA BENTHIC FORAMINIFERA Historically, benthic foraminifera have been studied more by paleontologists and geologists than by biologists, because of their abundant fossil record and economic use in petroleum exploration (e.g., Cushman, 1940). Paleontological study of deep-sea benthic foraminifera intensified with the beginning of scientific piston coring expeditions after World War II (e.g., Phleger et al., 1953), and of deep-sea drilling in the late 1960s (e.g., Berggren, 1972). At that time, little was known about the biology of foraminifera in general and deep-sea forms specifically, so that early studies consisted mainly of taxonomic descriptions and contained little interpretation of ecological or paleoceanographic information. Only recently, and specifically since such research programs as the Joint Global Ocean Flux Studies in the 1980s, have we learned more about the ecology of living foraminifera (e.g., Murray, 1991; Sen Gupta, 1999a; Smart, 2002; Gooday, 2003) and about the processes through which foraminifera calcify their tests (e.g., Hemleben et al., 1986; Hansen, 1999; Erez, 2003; Toyofuku and Kitazato, 2005). Even now, however, we are still ignorant of much of their biology and ecology, and paleoceanographic interpretation of deepsea benthic assemblages remains difficult (e.g., van der Zwaan et al., 1999; Murray, 2001; Jorissen et al., 2007). Benthic foraminifera are abundant deep-sea organisms, one of the principal eukaryote forms of life in the deep ocean, and constitute a large proportion of the eukaryotic deep-sea benthic biomass (e.g., Gooday, 1999, 2003). Like other deep-sea benthic organisms, they are locally highly diverse in normal marine environments, with more than 100 morphological species within relatively small sediment samples (e.g., Gooday, 1999; Gooday et al., 1998), but we do not know how this locally high diversity translates into regional diversity. As for other deep-sea biota, we do not truly understand why there is such high species richness among organisms many of which are deposit feeders that rely on organic detritus, in an environment where structural variety is

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apparently lacking in the endless tracts of sediment on the seafloor (Gage, 1996; Levin et al., 2001; Snelgrove and Smith, 2002; Rex et al., 2005). The role of environmental patchiness as described above is not yet well understood,, and neither is the contribution of lithoautotrophic prokaryotes to the overall oceanic carbon cycle and environmental patchiness on the seafloor (e.g., Dixon and Turley, 2001; Herndl et al., 2005). Since the 1970s, benthic foraminiferal assemblages have been used to derive information on the deep-sea environment of the geological past, with early papers interpreting Atlantic foraminiferal assemblages as reflecting the water-mass structure of that ocean (Streeter, 1973; Schnitker, 1974; Lohmann, 1978). Subsequent research looking for such linkages in different oceans and for different times was not entirely successful (e.g., Mackensen et al., 1995), and the development of transfer functions to derive quantitative expressions in which aspects of deep-sea benthic foraminiferal assemblages could be used as proxy for environmental parameters has proven difficult (van der Zwaan et al., 1999). This difficulty is probably explained by the fact that deepsea benthic foraminiferal assemblages are influenced by a combination of many parameters varying at different temporal and spatial scales and in many cases not independently from each other (e.g., Schnitker, 1994; Levin et al., 2001; Murray, 2001). Benthic foraminiferal proxies, many of which were reviewed by Gooday (2003) and Jorissen et al. (2007), include bathymetry (e.g., Hayward, 2004), organic matter flux, and oxygen concentrations in bottom and pore waters (e.g., Loubere, 1991, 1994, 1996; Kaiho, 1994b, 1999; Schmiedl et al., 1997, 2000); location and motion of redox fronts (with possibly related populations of Archaea and Bacteria) through the sediments (e.g., Fontanier et al., 2002, 2005); sediment type, temperature, bottom water chemistry (e.g., carbonate undersaturation, Bremer and Lohmann, 1982), hydrography (e.g., current flow; Schoenfeld, 2002), and hydrostatic pressure; and difficult-to-quantify parameters such as seasonality of the flux of organic matter, and relative amounts of labile and refractory organic matter (e.g., Gooday, 1988, 2002; Smart et al., 1994; Thomas et al., 1995; Loubere, 1998; Loubere and Fariduddin, 1999; Moodley et al., 2002; Fontanier et al., 2003). A strong correlation between benthic assemblages and one parameter is usually found only in extreme environments where organisms are strongly influenced by one limiting parameter, such as, for instance, severe dysoxia (low oxygen levels) to anoxia (lack of oxygen) (e.g., Bernhard, 1986; Sen Gupta and MachainCastillo, 1993; Bernhard et al., 1997; Moodley et al., 1998). Oxygen depletion resulting from organic enrichment is much more common along continental margins than on abyssal plains (e.g., Levin, 2003; Helly and Levin, 2004). Continental slopes and rises differ from abyssal plains because they are topographically complex and are more commonly subjected to vigorous currents and mass movements (e.g., turbidity currents, debris flows). Primary productivity along continental margins is higher overall than in open ocean, because of the prevalence of coastal upwelling in addition to nutrient discharges by rivers (e.g., Berger et al., 1988; Levin, 2003). On continental margins sedimentation rates

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are higher, sediments are more heterogeneous, and food particles may be supplied not only from primary productivity in the overlying waters but also by lateral transport of usually more refractory organic matter (e.g., Fontanier et al., 2005). Bottom- and pore-water oxygenation is usually inversely related to the flux of organic matter, because the oxidation of abundant organic matter causes the low oxygen conditions or even absence of oxygen that is common in areas of upwelling (e.g., Bernhard and Sen Gupta, 1999; Levin, 2003; Helly and Levin, 2004). Recent field and laboratory research focuses on these two inversely related parameters, the flux of organic matter (food) to the seafloor and the oxygen concentrations in bottom water and pore waters (e.g., Jorissen et al., 1995, 2007; Ohga and Kitazato, 1997; Schmiedl et al., 2000; Moodley et al., 1998; Gooday and Rathburn, 1999; Loubere and Fariduddin, 1999; van der Zwaan et al., 1999; Jorissen and Rohling, 2000; Moodley et al., 2000). It seems reasonable that food is an important limiting factor for deep-sea benthic foraminifera outside upwelling regions, especially on abyssal plains and along seamounts, because pelagic surface ecosystems are already severely nutrient limited if compared to, for example, terrestrial ecosystems (e.g., Tappan, 1986), and only a very small fraction of primary produced material reaches the bottom of the oceans (e.g., Smith et al., 1997). The interplay of the two interrelated factors, oxygen availability and food supply, and their effects on the benthic faunal assemblage and its position within the sediments was discussed by Jorissen et al. (1995) in the “TROX-model,” and modified and refined by various authors (Jorissen et al., 1998; Jorissen, 1999; Fontanier et al., 2002; and Gooday, 2003). In this model, microhabitat occupancy (the places where foraminifera live within the sediment) is correlated to the availability of food and oxygen: in a continuum of increasing food supply from oligotrophic through mesotrophic to eutrophic, oxygen levels are negatively correlated to food supply. At the oligotrophic extreme, populations are mainly limited by food, with a “critical food level” at shallow depths (a few centimeters). At the eutrophic extreme, populations are limited by oxygen levels, with a critical oxygen level finally reaching the sediment-water interface or even moving into the water column when bottom waters become anoxic. In oligotrophic settings (e.g., abyssal plains) foraminifera are highly concentrated in the uppermost levels of the sediment: most food particles are used up by organisms dwelling close to the sediment-water interface, very little organic matter is buried in the sediment except for that worked down by bioturbating organisms such as echiuran worms, and pore waters are well oxygenated. As the food supply increases, epifaunal to shallow infaunal forms use only part of the food. The remainder of the food is buried to greater depths, so that species can live and feed in the sediment, at depths ranging from epifaunal to shallow infaunal (0–1.5 cm), to intermediate infaunal (1.5–5.0 cm), to deep infaunal (5–10 cm). The deepest level of occurrence is limited either by the food available (toward the oligotrophic part of the continuum) or by the low oxygen levels (toward the eutrophic end of the continuum). The highest species diversity occurs in

mesotrophic regions, with co-occurrence of epifaunal through deep infaunal forms (Fig. 1 in Gooday, 2003). Epifaunal and infaunal groups may differ in, and thus be recognized by, their overall test morphology (e.g., Corliss, 1985; Corliss and Chen, 1988; Thomas, 1990a; Kaiho, 1991), but there is some confusion about the use of the terms “epifaunal” and “infaunal.” The above authors used the term “epifaunal” for foraminiferal species living on the surface and in the uppermost 0–1 cm sediment. In soft sediments, the sediment-water interface is not sharply defined, and living exactly at this interface is difficult to impossible. Buzas et al. (1993) pointed out that species within the top 1 cm of sediment are actually living within the sediment because many foraminifera are much smaller than 1 cm, and thus should be called “shallow infaunal.” Later publications tend to not use the word “epifaunal,” unless indicating species such as Cibicidoides wuellerstorfi, which prefer to live on objects sticking out above the sediment-water interface (e.g., Fig. 3 in Altenbach and Sarnthein, 1989). In general, benthic foraminifera with plano-convex, biconvex, and rounded trochospiral, tubular, and coiled-flattened tests have been observed be epifaunal to shallow infaunal. Foraminifera living in the deeper layers of the sediment have cylindrical or flattened tapered, spherical, rounded planispiral, flattened ovoid, globular unilocular, or elongate multilocular tests. For many taxa, however, the relation between test morphology and microhabitat has not been directly observed but is extrapolated from data on other taxa (e.g., Jorissen, 1999). In one of the few studies evaluating the correlation between test morphology and microhabitat statistically, such assignments for modern foraminifera were shown to be accurate only ~75% of the time (Buzas et al., 1993). A matter of debate is the importance of the food flux versus that of oxygenation in determining the foraminiferal assemblages, as reviewed most recently by Gooday (2003). Most authors agree that in generally oxygenated conditions (i.e., oxygen concentrations above ~1mg/L) food is the more important determinant (e.g., Rathburn and Corliss, 1994; Morigi et al., 2001), whereas Kaiho (1994b, 1999) places more importance on oxygen levels in bottom waters. Several authors (as reviewed in Gooday, 2003) agree that the boundary between more oligotrophic and more eutrophic regions is at about a flux level of 2–3 g Corg m–2yr–1. The generally more eutrophic continental margins contain assemblages with more abundant infaunal species, including species belonging to the genera Bolivina, Bulimina and Uvigerina, and these species are considered indicative of eutrophic conditions. In detail, and on short temporal and small spatial scales, the situation is considerably more complex than given in the TROX model (e.g., Linke and Lutze, 1993). Different foraminiferal species have different food preferences (e.g., Lee, 1980; Goldstein and Corliss, 1994; Heinz et al., 2002; Suhr et al., 2003). Species react differently to food pulses, with some species reacting rapidly and opportunistically by fast reproduction (e.g., Altenbach, 1992; Linke et al., 1995; Ohga and Kitazato, 1997; Moodley et al., 2000, 2002), especially species that feed on fresh phytodetritus (Gooday, 1988, 1993; Moodley et al., 2002; Suhr et al., 2003),

Cenozoic mass extinctions in the deep sea whereas other species grow more slowly (Nomaki et al., 2005). Many species do not permanently live at the same depth below the sediment-water interface, but move vertically through the sediments (e.g., Kaminski et al., 1988; Linke and Lutze, 1993; Bornmalm et al., 1997; Ohga and Kitazato, 1997; Gross, 1998; Gooday and Rathburn, 1999; Jorissen, 1999; Gross, 2000; Fontanier et al., 2002; Geslin et al., 2004), either reacting directly to food pulses (Altenbach, 1992; Linke et al., 1995; Heinz et al., 2002), or following pore-water oxygen gradients, which define different redox levels characterized by specific bacterial-archaeal populations (Moodley et al., 1998; Gross, 1998, 2000; Geslin et al., 2004). Benthic foraminiferal proxies have been used extensively in an attempt to reconstruct export productivity (e.g., Altenbach et al., 1999). Proxies for primary productivity include the Benthic Foraminiferal Accumulation Rate (BFAR; Herguera and Berger, 1991) and the more statistically complex method of Loubere (1994, 1996). The approach appears to give good results mainly in well-oxygenated sediments where no carbonate dissolution occurs, but the validity of the quantitative calculation may be questionable when there is variation in the type of organic matter deposited (e.g., Guichard et al., 1997). The correlation between BFAR and export productivity may not be linear in the presence of opportunistically blooming, phytodetritus-exploiting species (Schmiedl and Mackensen, 1997). The fact that in many regions there is a significant correlation between benthic foraminiferal accumulation rate and primary productivity in surface waters, however, indicates that present deep-sea faunas receive most of their food (directly or indirectly) from photosynthetic primary producers. Using benthic foraminifera to reconstruct export productivity can be problematic because such a proxy must be calibrated in the present oceans, where the correlation between export productivity and food arriving at the seafloor is not well known quantitatively. The single-celled algae (including prokaryotes, Kolber et al., 2000) at the base of the pelagic food chain are not efficiently deposited to the seafloor as single particles. For efficient transfer, material must be aggregated in larger particles (marine snow; e.g., Turley, 2002), in sticky, seasonally produced phytodetritus (e.g., Jackson, 2001; Beaulieu, 2002) in which the “stickiness” increases by the exudation by phytoplankton of polysaccharides in Transparent Exopolymer Particles (TEP; Engel et al., 2004), in diatom mat aggregates (e.g., Kemp et al., 1995, 2000), in fecal pellets of zooplankton or nekton, ballasted by siliceous and carbonate tests (e.g., François et al., 2002; Klaas and Archer, 2002) or by terrigenous dust (Ittekkot, 1993), or in tunicate feeding structures (Robison et al., 2005). Such aggregated food reaches the deep-sea floor in a few weeks and contains the labile, fresh organic material that is used preferentially by some benthic foraminifera (e.g., Gooday, 1988, 1993, 2002; Ohga and Kitazato, 1997; Suhr et al., 2003). Only a very small fraction of the primary produced biomass reaches the seafloor (0.01%–1.0%; e.g., Murray et al., 1996), and there is no linear correlation between productivity and flux below 2000 m at high productivities (>200 gC m–2yr–1) (Lampitt and Antia, 1997). There

5

is a discrepancy between measured fluxes of sinking particulate organic matter and food demand (sediment community oxygen consumption), with the fauna apparently consuming more food than is supplied, as measured in the Pacific Ocean (Smith and Kaufmann, 1999; Smith et al., 2002) and as averaged over the world’s oceans (Del Giorgio and Duarte, 2006). Some authors argue that high seasonality results in high efficiency of transport to the seafloor (Berger and Wefer, 1990), whereas others argue the reverse (François et al., 2002). In addition, there is no agreement whether carbonate or siliceous tests are more efficient in ballasting organic matter (compare François et al., 2002, and Klaas and Archer, 2002, with Katz et al., 2005), and whether dust, carbonate, or silica tests function as ballast for organic matter at all (Passow, 2004). Increased supply of calcareous tests intuitively appears to lead to increased ballasting, but investigations of Emiliania huxleyi blooms under high atmospheric pCO2 levels have documented increased transport of organic matter to the seafloor with less calcification, because the production of the sticky TEPs increases under such conditions (Delille et al., 2005). Locally, organic matter may be supplied mainly by lateral transport (e.g.,, Fontanier et al., 2005). Lithoautotrophic organic matter, which is produced in situ on the seafloor or in overlying waters below the photic zone (e.g., Karner et al., 2001; Herndl et al., 2005; Bach et al., 2006), could add to the overall food supply even outside the direct environment of hydrothermal vents or cold seeps. Bacterial oxidation of methane in hydrothermal plumes in the present oceans contributes an amount of organic carbon up to 150% that of the surface-produced organic matter reaching the depth of the plume (2200 m) in the northeast Pacific (Roth and Dymond, 1989; de Angelis et al., 1993). In view of these uncertainties about food transfer to the deep-sea in the present oceans, we are uncertain regarding the quantitative correlation between present foraminiferal parameters and primary productivity in the photic zone. Because only such a minute fraction of primary produced organic matter reaches the seafloor, relatively small changes in efficiency of this transfer may have a major impact on the amount of food reaching benthic assemblages, even at constant productivity (see also Katz et al., 2005). CENOZOIC BENTHIC FORAMINIFERAL FAUNAS In contrast to planktic foraminifera, deep-sea benthic foraminiferal species have, on average, long species lives, and thus cannot be used for detailed stratigraphic subdivision of geological time: they are not good “guide fossils” (e.g., Boltovskoy, 1980, 1987; van Morkhoven et al., 1986; Tappan and Loeblich, 1988; Thomas, 1992a; Culver, 1993). The majority of modern calcareous smaller benthic foraminifera in the deep oceans belong to the orders Rotalida and Buliminida (Sen Gupta, 1999b), which became common in the deep oceans gradually, after the Cenomanian-Turonian Oceanic Anoxic Event, with many common genera present from the Campanian (Kaiho, 1994a, 1998).

6

Thomas

These taxa thus became common in the deep oceans during the later part of the “Mesozoic Marine Revolution,” a time of reorganization of ecosystems and evolution of modern forms of many marine animals and eukaryotic plankton (e.g., Vermeij, 1977; Bambach, 1993; Katz et al., 2005). The deep-sea benthic foraminifera of the later Campanian and the Maastrichtian contain many components that persist in Paleocene faunas, including common species such as Nuttallides truempyi and Stensioeina beccariiformis (e.g., Cushman, 1946; van Morkhoven et al., 1986; Thomas, 1990b, 1992a; Kaiho, 1994a, 1998; Alegret and Thomas, 2001). Benthic foraminifera did not suffer significant extinction across the Cretaceous- Paleogene boundary (review by Culver, 2003): their most severe extinction of the Cenozoic occurred at the end of the Paleocene (review by Thomas, 1998). Assemblage zones for Cenozoic bathyal and abyssal faunas have been recognized for the southern oceans (Thomas, 1990a), for the Indian Ocean (Nomura, 1991, 1995), and for the global oceans (Berggren and Miller, 1989). These zones can be simplified to four Cenozoic faunas (Miller et al., 1992; Thomas, 1992a) (Fig. 1): (1) a “Cretaceous Fauna,” which survived from the Late Cretaceous and suffered abrupt extinction at the end of the Paleocene; (2) an early-middle Eocene “Paleogene fauna,” which underwent a gradual but severe turnover (E/O) through the late Eocene and earliest Oligocene (Corliss, 1981; Thomas, 1992a,

δ18O 5

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1992b; Thomas and Gooday, 1996); (3) an Oligocene–early Miocene “Transitional Fauna,” which underwent a gradual turnover (mMIO) in the middle Miocene (Woodruff and Douglas, 1981; Thomas, 1986; Woodruff and Savin, 1989); and (4) the “Modern Fauna.” An additional transition (mPL) was recognized in the middle Pleistocene (ca. 1.2–0.6 Ma), with the extinction of many cylindrical species having complex (e.g., dentate, cribrate, lunate) apertural shapes (Weinholz and Lutze, 1989; Schoenfeld, 1996; Hayward, 2001, 2002; Kawagata et al., 2005). The three gradual benthic foraminiferal faunal turnovers (E/O, mMIO, mPL) are similar in paleoceanographic setting: the first two occurred during periods of global cooling and growth of polar ice sheets (e.g., Zachos et al., 2001), and the last one during the period of intensification of Northern Hemisphere glaciation (Zachos et al., 2001; Tziperman and Gildor, 2003). During all three turnovers there was a loss of “cylindrical” species, mainly uniserial or elongate biserial, belonging to the stilostomellids, pleurostomellids, or uniserial lagenids, some of the latter group surviving to today (e.g., Thomas, 1986; Hayward, 2001, 2002; Kawagata et al., 2005). Many of these cylindrical species (including all stilostomellids) became extinct during the Mid Pleistocene Revolution (so that we have no information on their ecology), after decreasing in abundance during the Eocene-Oligocene and middle Miocene turnovers (Fig. 2). These cylindrical species

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70 0° 4° 8° 12° Ice-free Temperature (°C)

Figure 1. Cenozoic deep-sea benthic foraminiferal faunas (modified after Thomas et al., 2000) as compared with the deep-sea carbon and oxygen isotopic record (Zachos et al., 2001). *mPL— Mid-Pleistocene benthic foraminiferal turnover, also known as the Stilostomella extinction (Hayward, 2001). P/E—Paleocene/Eocene benthic foraminiferal extinction.

Cenozoic mass extinctions in the deep sea

uniserial lagenids, %

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Figure 2. Relative abundance of cylindrical taxa during the Cenozoic. The vertical blue bars indicate the periods of faunal turnover (Fig. 1). P/E—Paleocene/Eocene benthic foraminiferal extinction. E/O— Eocene/Oligocene faunal turnover. mMIO—Middle Miocene faunal turnover. mPL—Middle Pleistocene faunal turnover. Data are compiled from DSDP/ODP Sites in the equatorial Pacific, with ages recalculated according to Berggren et al., 1995. Red—Site 865 (Thomas, 1998, unpublished data); Sites 573, 574, 575 (Thomas, 1985). Blue— Weddell Sea Sites 689, 690 (Thomas, 1990a). Green—North Atlantic Sites 608, 610 (Thomas, 1987; see also Thomas and Gooday, 1996). Note that only some groups of the uniserial lagenids became extinct in the middle Pleistocene, with several surviving until today.

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may have lived infaunally and reflect a relatively high food supply: their abundances correlate positively with those of genera that in present oceans are indicative of high food and/or low oxygen (e.g., Bulimina, Uvigerina, Bolivina) in the North Atlantic (Kawagata et al., 2005). There are biogeographical differences between the various cylindrical taxa (Fig. 2). Uniserial lagenids were more abundant at low latitudes and middle-upper bathyal depths, and decreased in abundance strongly in the late Eocene, whereas the pleurostomellids did so mainly in the middle Miocene. Stilostomellids were more abundant at high latitudes, especially in the late Eocene, and decreased in both the Eocene-Oligocene and the middle Miocene turnovers. As a result of these three turnovers, Recent open-ocean deepsea faunas (away from continental margins) differ strongly from middle Eocene and older ones (e.g., Boltovskoy, 1984, 1987; Miller et al., 1992; Thomas et al., 2000), which lived in “greenhouse” oceans that were 10–12 °C warmer than the present oceans (e.g., Zachos et al., 2001). The Eocene and older faunas living remote from the continents resemble present continental-margin faunas in common morphotypes: high percentages of high food/low oxygen indicator genera, cylindrical taxa, and taxa in the Order Buliminida (Thomas et al., 2000). These “greenhouse” faunas contained only very rare phytodetritus-using species (e.g., Epistominella exigua), which are common in present open-ocean settings (e.g., Gooday, 2003). Such phytodetritus-using species bloom opportunistically when fresh, labile organic material reaches the seafloor, and they rapidly increased in abundance during the E/O turnover. In the middle Miocene they became common even in the equatorial Pacific (Thomas, 1985; Thomas and Gooday, 1996; Thomas et al., 2000). Miliolid taxa earlier had become common at neritic depths, and migrated into the deep oceans during the middle Miocene turnover (Thomas, 1986; 1992a), when the suspensionfeeding Cibicidoides wuellerstorfi evolved, with a first appearance slightly earlier in the Pacific than in the Atlantic (Thomas and Vincent, 1987). At greater depth and at high latitudes (close to the lysocline), the relative abundance of Nuttallides umbonifera (a species seen as indicative of corrosive bottom waters and/or oligotrophy) strongly increased in abundance during the E/O turnover (Thomas, 1992a; Thomas et al., 2000). The modern assemblages typical for seasonal delivery of fresh phytodetritus to the seafloor had no early Paleogene counterpart: bentho-pelagic coupling, in which labile organic matter is supplied in seasonal pulses, may have originated or intensified with the establishment of the Antarctic ice sheet in the earliest Oligocene (Zachos et al., 2001), when stratification of the oceans increased (e.g., Schmidt et al., 2004), seasonality of productivity increased (e.g., Thomas and Gooday, 1996), the importance of diatoms as primary producers increased (Katz et al., 2004), the size of diatoms increased (Finkel et al., 2005), and fresh phytodetritus (arriving at the seafloor only weeks after having been produced) became an important part of the food delivered to the deep-sea benthos. Early Paleogene deep-sea benthic foraminiferal assemblages may thus have differed from modern ones by not having the niche of

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“phytodetritus species” filled. The Paleogene assemblages contained mainly deposit feeders (as defined by Goldstein and Corliss, 1994), with different niches defined by their location with regard to redox fronts, the presence or absence of xenophyophoran tests, and other differences at a small spatial scale. In this view, the E/O turnover represents an important, structural change in deep-sea faunas, with modern-type assemblages becoming established (Boltovskoy, 1984, 1987; Thomas et al., 2000). The Eocene/Oligocene, middle Miocene, and middle Pleistocene benthic foraminiferal faunal turnovers occurred during periods of global cooling, and represent subsequent but similar steps during which the open-ocean faunas gradually took on their modern composition, losing various groups that we consider to indicate either a high food supply or low oxygenation (Kaiho, 1991; Kawagata et al., 2005), or a less seasonally pulsed overall food supply (Thomas and Gooday, 1996; Thomas et al., 2000; Thomas, 2003). In these speculations it should not be forgotten that morphotype assignments are problematic in the present oceans (Buzas et al., 1993), and thus even more so for extinct species. The apparent loss of high-food species during global cooling presents a paleoceanographic problem, because cooling, increasingly vigorous ocean circulation, and upwelling, possibly in addition to more vigorous chemical weathering and delivery of nutrients to the oceans, have been thought to induce increased oceanic productivity, (e.g., Brasier, 1995a, 1995b; Katz et al., 2004; Finkel et al., 2005). The problem is exacerbated because in the early Paleogene deep-water temperatures were ~10°C higher than today, so that metabolic rates of foraminifera were higher by as much as a factor of 2 (e.g., Hallock et al., 1991; Gillooly et al., 2001), thus requiring twice as much food in order to keep the same faunal structure. Research on modern faunas does not support the hypothesis (Kaiho, 1991) that the faunal change in the late Eocene–Oligocene resulted mainly from increased oxygenation, because benthic foraminifera do not appear to be influenced by oxygenation at the levels reconstructed for these times (e.g., Gooday, 2003). It seems improbable that export productivity in the middle Eocene and earlier was more than twice as high as at present (e.g., Brasier, 1995a, 1995b), but the efficiency of food transfer to the seafloor (as discussed above for the present oceans) may very well have differed in the “greenhouse” oceans, as suggested by the proposal that a larger fraction of organic matter was preserved in Paleogene sediments (Kump and Arthur, 1997). In view of our lack of understanding of food transfer in the present oceans, we cannot be certain about such processes in the “greenhouse” oceans. Ecosystem modeling does not clearly answer whether warming would result in higher or lower net global productivity (Sarmiento et al., 2004), or export productivity (Laws et al., 2000). Food-web structure would probably be different in a warmer world (e.g., Petchey et al., 1999) without polar sea ice (Loeb et al., 1997), and the effects of such changes in ecosystems structure are not understood. We can speculate on several, not mutually exclusive, possibilities (see also Thomas et al., 2000; Thomas, 2003):

1. The ocean circulation may have been not just quantitatively but also qualitatively different from that in the present oceans, with ‘greenhouse’ oceans dominated by eddy rather than gyral circulation (Hay et al., 2005), or increased hurricane activity and vertical mixing, resulting in increased poleward heat transport (Emanuel, 2001, 2002). Such different circulation patterns might have resulted in more efficient transfer of food to the seafloor as the result of more extensive vertical water motion and more vigorous deep vertical mixing. Emanuel’s (2001) discussion argues for increased openocean upwelling, thus less difference in primary productivity between continental margins and open ocean, as supported by the benthic foraminiferal evidence. 2. Higher temperatures of the oceans could have resulted in lower oxygenation, thus in less degradation of organic matter. This does not appear to be probable because there is no significant correlation between oxygenation and organic matter preservation in the present oceans (e.g., Hedges and Keil, 1995; Kristensen et al., 1995). In addition, present deep-sea faunas in relatively warm oceans (e.g., the Red Sea) are oligotrophic in character (e.g., Thiel et al., 1987; Edelman-Furstenberg et al., 2001). 3. Higher temperatures would result in higher metabolic rates for bacteria, which determine organic matter degradation and transformation, as well as lithoautotrophic production. More active bacteria could have resulted in a more active “bacterial loop,” more conversion of dissolved organic carbon into particulate organic carbon, thus enhanced food supply for zooplankton, delivering more food (fecal pellets) to the seafloor (e.g., Jannasch, 1994; Verity et al., 2002). Alternatively, higher bacterial activity could mean higher use of refractory carbon, and increase in bacterial biomass that can be taken in by benthic foraminifera. In addition, increasing temperatures lead to the dominance of different bacterial groups, producing different organic compounds (Weston and Joye, 2005), for which different groups of foraminifera might have a preference. 4. Dominant “greenhouse” ocean primary producers differed from those in the present “icehouse” ocean: diatoms strongly increased in importance and size after the end of the Eocene (e.g., Katz et al., 2004; Finkel et al., 2005). Many diatoms produce mucus that causes coagulation of organic matter; thus one would expect more diatom productivity to result in more rapid transfer of organic matter to the seafloor. François et al. (2002) and Klaas and Archer (2002), however, argue that transfer is less efficient in diatom-dominated systems than in carbonate (calcareous nannofossil) dominated systems. If this was true for the early Paleogene, food transfer could have been more efficient due the greater prevalence of calcareous primary producers. 5. Under the higher pCO2 levels that may have been present in the atmosphere in the Paleogene (Zachos et al., 2001; Pagani et al., 2005) calcareous nannoplankton may have calcified less in the more acid oceans (Feely et al., 2004; Orr et al., 2005), unexpectedly leading not to decreased deposition of organic

Cenozoic mass extinctions in the deep sea matter because of decreased carbonate ballasting, but to increased exudation of sticky polysaccharides and increased deposition of organic carbon (Delille et al., 2005). 6. In the Paleogene, the contribution to the seafloor food supply by in situ lithoautotrophic productivity was larger (Thomas, 2003). Benthic foraminifera living in cold-seep areas are not taxonomically different from high-food/low oxygen species (Rathburn et al., 2000; Bernhard et al., 2000, 2001; Barbieri and Panieri, 2004), and we thus cannot distinguish overall high food supply from supply by chemosynthesis. The higher temperatures could have speeded up metabolism of bacteria, resulting in increased methane production at the lower oxygenation, possibly in combination with higher delivery of organic matter to the seafloor. They could also have led to production of more labile compounds by bacteria, which foraminifera can take up more easily (Weston and Joye, 2005). 7. The hypothesis that benthic foraminiferal assemblages indicate a higher overall food supply to the seafloor is incorrect: what is really indicated is a less seasonal delivery of organic material in the early Paleogene (e.g., Thomas and Gooday, 1996; Ohkushi et al., 1999). The abundant occurrence of phytodetritus species apparently does not indicate overall levels of primary productivity, but rather a strongly seasonal delivery of food (e.g., Gooday, 2003). Such a strongly seasonal food supply can be used only by species that can rapidly and opportunistically react, possibly causing the gradual demise of species that may have been specialized detritus feeders, with their complex apertural structures directing pseudopodial flow. BENTHIC FORAMINIFERA AT THE CRETACEOUS-PALEOGENE BOUNDARY Most earth scientists accept that the end-Cretaceous mass extinction was caused at least in part by the impact of a meteorite with a diameter of ~10 km (Alvarez et al., 1980) on the northern Yucatán peninsula (Hildebrand et al., 1991). Deep-sea benthic foraminifera are among the survivors. Until recently, net extinction of benthic foraminifera was alleged to have been more severe in shallower waters (e.g., Thomas, 1990b; Kaiho, 1992, 1994a; Coccioni and Galeotti, 1998), but the excellent review by Culver (2003) documents that shallow-dwelling species were not more severely affected than deeper-dwelling ones. Benthic foraminiferal assemblages underwent temporary changes in community structure in calcareous species (Culver, 2003) as well as in agglutinated taxa (e.g., Kuhnt and Kaminski, 1996; Culver, 2003; Bak, 2005; Kaminski and Gradstein, 2005). These changes in community structure varied geographically and bathymetrically (e.g., Coccioni and Galeotti, 1998; Alegret and Thomas, 2005) but did not result in significant net extinction. At locations relatively close to the impact location, e.g., in Mexican sections (Alegret and Thomas, 2001; Alegret et al., 2001) and in the northwestern Atlantic (Alegret and Thomas, 2004), the record is interrupted by slump-beds containing alloch-

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thonous neritic foraminifera, so that detailed faunal records across the boundary cannot be obtained. At other locations, e.g., in the Caravaca and Agost sections in Spain (Coccioni et al., 1993; Coccioni and Galeotti, 1994; Kaiho et al., 1999; Alegret et al., 2003), dark clay layers contain low-diversity, high food/ low oxygen assemblages, indicative of low oxygen conditions. In the equatorial Pacific, assemblages and benthic foraminiferal accumulation rates indicate a short-time increase in food delivery, but evidence for low oxygen conditions is weak (Alegret and Thomas, 2005). Recovery of faunas, i.e., of diversity and abundance of infaunal taxa, varied at different locations, from 100 to 300 k.y. (Alegret and Thomas, 2005). This paper will not address the regionally variable patterns of assemblage composition after the Cretaceous-Paleogene boundary, except to mention that there is no agreement on the detailed environmental interpretation of the generally low-diversity assemblages occurring just after that boundary (see, e.g., discussions on the El Kef section by Speijer and van der Zwaan, 1996, and by Culver, 2003). Most authors agree that the temporary faunal restructuring of benthic foraminifera was caused by the collapse of the pelagic food web (e.g., Thomas, 1990a, 1990b; Widmark and Malmgren, 1992; Coccioni et al., 1993; Kuhnt and Kaminski, 1996; Speijer and van der Zwaan, 1996; Peryt et al., 1997, 2002; Culver, 2003; Alegret et al., 2001, 2003, 2004; Alegret and Thomas, 2001, 2004, 2005). It is difficult to understand, however, how such relatively minor and reversible changes in benthic faunal assemblages could have been the response to a major, long-term collapse of oceanic productivity, with biological productivity low for hundreds of thousands to several millions of years after the asteroid impact, and the slow-down or even stop of the “biological pump” of organic matter to the seafloor (“Strangelove Ocean”). Evidence for such a “Strangelove Ocean” consists of the collapse of the gradient between benthic and planktic (foraminiferal and/or bulk carbonate) carbon isotope values (e.g., Arthur et al., 1979; Hsü et al., 1982; Hsü and McKenzie, 1985; Zachos and Arthur, 1986; Zachos et al., 1989). The end-Cretaceous benthic foraminiferal assemblages may have been characterized by less intense bentho-pelagic coupling than present faunas (as argued above; Thomas et al., 2000), but even under such conditions they should have shown a more severe change in community structure than observed, if food supplies remained so extremely low for millions of years. At times of less intense bentho-pelagic coupling, most benthic foraminifera probably adopted some form of depositfeeding lifestyle, and such a lifestyle has been argued to be a possible exaptation to survive an impact-driven productivity crash. Recent evidence does not support this hypothesis, however (Levinton, 1996; Jablonski, 2005), even though the wide geographic range of benthic foraminiferal genera might have predisposed them for survival (Jablonski, 2005) More recently, it was proposed that productivity (in terms of biomass) recovered as soon as light returned after the impact, although the plankton diversity remained low and the transfer of organic matter to the seafloor remained limited (D’Hondt et al.,

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1998; D’Hondt, 2005; Coxall et al., 2006). During this partial recovery, lasting several hundred thousand years, lower gradients of benthic-planktic carbon isotope values persisted because of a severe decrease in food transfer to the seafloor due to ecosystem reorganization (e.g., loss of fecal pellet producers as a result of the mass extinction; shift to smaller-celled primary producers) (D’Hondt et al., 1998; Adams et al., 2004). The lack of significant extinction of benthic foraminifera does not agree with the “Strangelove Ocean” model, and favors instead a model in which productivity as well as food transfer to the seafloor recovered more quickly, and more fully, than proposed even in the “living ocean” model (D’Hondt et al., 1998; Adams et al., 2004; D’Hondt, 2005; Coxall et al., 2006). Calcareous nannoplankton suffered very high rates of extinction, but other primary producers such as diatoms did not (Kitchell et al., 1986), cyanobacteria may not have been affected (D’Hondt et al., 1998), and the dinoflagellate calcareous cyst Thoracosphaera bloomed opportunistically worldwide (e.g., Thierstein, 1981; Perch-Nielsen et al., 1982; Gardin and Monechi, 1998). Surviving phytoplankton could be expected to bloom as soon as light conditions allowed, because the extinction of calcareous nannoplankton lessened competition for nutrients. The blooms of opportunistic phytoplankton may not have been global, but they occurred locally or regionally and led to local or regional anoxia, as observed, for instance, in the Caravaca and Agost sections of southern Spain (Coccioni et al., 1993; Coccioni and Galeotti, 1994; Kaiho et al., 1999; Alegret et al., 2003). There is not enough evidence to fully evaluate the severity, geographic and depth extent of hypoxia/anoxia after the Cretaceous-Paleogene boundary, but the lack of benthic foraminiferal extinction contradicts evidence in favor of global anoxia (e.g., Kajiwara and Kaiho, 1992). Transport of organic matter to the seafloor may have recovered faster than envisaged by D’Hondt et al., (1998). Coagulation of organic particles by sticky diatoms and cyanobacteria may have assisted in forming large particles for rapid deposition (see above; Jackson, 2001; Armstrong et al., 2001), and various methods of ballasting particles with biogenic silica or terrigenous dust may have remained effective, even with less biogenic carbonate available. If atmospheric pCO2 levels were very high after the impact (Beerling et al., 2002), calcification of the few surviving calcareous nannofossils may have decreased, but decreased calcification may have led to increased delivery of organic matter to the seafloor because of increased formation of sticky polysaccharides (Delille et al., 2005; Engel et al., 2004). If both productivity and food transfer to the deep seafloor recovered faster than previously assumed, the recovery of marine ecosystems would be similar to the rapid recovery postulated for terrestrial ecosystems (Beerling et al., 2001; Lomax et al., 2001). The lack of extinction of benthic foraminifera could then be understood, but the persistent collapse of benthic-planktic carbon isotope gradients must be explained. I suggest that this collapse may not necessarily reflect a lack of operation of the biotic pump. As a first possibility, the lighter values in bulk carbonate and planktic foraminiferal tests, reflecting the carbon isotope values

of total dissolved carbon in surface waters, may not reflect a drop in productivity. A negative carbon isotope anomaly has also been observed in terrestrial materials, indicating that a marine-productivity explanation is not sufficient (Arinobu et al., 1999; Arens and Jahren, 2000). The negative carbon isotope anomaly might have been caused by an input of light carbon in the surface ocean-atmosphere system (not penetrating into the deep sea), either as the result of biomass burning (Ivany and Salawitch, 1993) or methane liberation from dissociation of gas hydrates due to massive slumping on continental margins (Max et al., 1999; see also discussion in Alegret et al., 2003; Norris and Berger, 2003; Day and Maslin, 2005). Alternatively, at least part of the surface isotope signal may reflect “vital effects” (e.g., Rohling and Cooke, 1999; Stoll and Ziveri, 2002; Maslin and Swann, 2005; Ziveri et al., 2003). The carbon isotope values reflecting isotope values of total dissolved carbon in surface waters must by necessity be measured on calcareous nannofossils (bulk records) and/or planktic foraminifera. Both groups underwent severe extinction, so that post- and preextinction records are derived from different species than the pre-extinction records. Post-extinction calcareous nannoplankton is dominated by bloom species such as Thoracosphaera, Braarudosphaera and Biscutum (Perch-Nielsen et al., 1982). Some of these taxa (Thoracosphaera) are calcareous dinoflagellates, and Cretaceous (Friedrich and Meijer, 2003) as well as Recent (Zonneveld, 2004) species of calcareous dinoflagellates have very light carbon isotope signatures. Finally, the bulk record at some sites may be affected by diagenesis, which is common in low-carbonate sediments (e.g., Zachos et al., 2005). These three possibilities are not mutually exclusive, and the Cretaceous-Paleogene surface-bottom carbon isotope gradient collapse may reflect a more complex signal than one of collapsed productivity only. PALEOCENE/EOCENE BENTHIC FORAMINIFERAL EXTINCTION It has long been known that a major extinction of deep-sea benthic foraminifera occurred at the end of the Paleocene. Cushman (1946) placed the Cretaceous-Paleogene boundary at the end of the Paleocene, because the total foraminifera (planktic + benthic) show a much larger species turnover at that time, being dominated by the numerous benthic species (though low numbers of specimens). The extinction was documented in Trinidad by Beckmann (1960), in Austria by von Hillebrandt (1962), and in Italy by Di Napoli Alleata et al. (1970) and Braga et al. (1975) (review by Thomas, 1998). The extinction was first described in detail by Tjalsma and Lohmann (1983), using data from drill holes in the Atlantic Ocean and Gulf of Mexico. The scope and rapidity of the extinction, however, was not realized in these earlier studies, mainly because detailed, high-resolution age models were lacking. The event was first described as a major, rapid extinction (duration <10 k.y.) by Thomas (1989, 1990a). It was coeval with an episode of extreme global warming now called the Paleocene-

Cenozoic mass extinctions in the deep sea Eocene Thermal Maximum (PETM), with temperature increases of up to 9–10 °C in high-latitude sea surface temperatures, ~4–5 °C in the deep sea and in surface waters in equatorial regions (Zachos et al., 2003; Tripati and Elderfield, 2004, 2005), ~5 °C on land at midlatitudes in continental interiors (e.g., Wing et al., 2005), and in the Artic Ocean (Sluijs et al., 2006). During this time, humidity and precipitation were high, especially at middle to high latitudes (e.g., Bowen et al., 2004, 2006; Pagani et al., 2006). Diversity and distribution of surface marine and terrestrial faunas and floras shifted, with rapid migration of thermophilic biota to high latitudes, as well as rapid evolutionary turnover (e.g., Crouch et al., 2001; Wing et al., 2005; papers in Wing et al., eds., 2003). Deep-sea benthic foraminifera, in contrast, suffered severe extinction (30%–50% of species), although these organisms survived such environmental crises as the asteroid impact at the end of the Cretaceous without significant extinction (e.g., Thomas, 1989, 1990a, 1990b, 1998; Thomas et al., 2006; Culver, 2003). Severe dissolution occurred in many parts of the oceans (Thomas, 1998), with the calcium carbonate compensation depth (CCD) shifting upward by at least 2 km in the southeastern Atlantic (D. Thomas et al., 1999; Zachos et al., 2005), although less in the Pacific Ocean (Colosimo et al., 2005). Carbon isotope data on planktic and benthic foraminiferal tests and bulk marine carbonates (first documented by Kennett and Stott, 1991; Thomas and Shackleton, 1996), and on soil organic matter, soil carbonates, and herbivore teeth (Koch et al., 1992, 2003), show that there was a large perturbation to the global carbon cycle, affecting the whole ocean-atmosphere system, as seen by a negative excursion (carbon isotope excursion, CIE) of at least 2.5‰ in oceanic records (e.g., Zachos et al., 2001), and 5–6‰ in terrestrial records (e.g., Magioncalda et al., 2004; Koch et al., 2003; Bowen et al., 2004). The isotope anomalies indicate a rapid onset of these anomalies (<<20 k.y.; Roehl et al., 2000), followed by return to more normal values on time scales of 105 years (Westerhold et al., 2007; Sluijs et al., 2007). Researchers agree that the episode of rapid global warming was caused by the massive input of isotopically light carbon into the ocean-atmosphere system. Although there has been intensive research and vigorous discussion, there is no agreement on the source of the added carbon (e.g., Sluijs et al., 2007). The usual mechanisms called upon for carbon isotope excursions do not work at the PETM time scale: the isotopic composition of volcanic emissions is not light enough and their rate of emission too slow, the weathering of organic carbon-rich rocks is too slow, and the CIE is too large to have been caused by destruction of land biomass (e.g., Thomas and Shackleton 1996). Since 1995 (Dickens et al., Matsumoto) the most widely accepted hypothesis for the cause of the isotope anomalies has been the release of ~2000–2500 Gt of isotopically very light (~-60‰) carbon from methane clathrates in oceanic reservoirs, with subsequent severe greenhouse-gas induced warming. Oxidation of the methane in the oceans could have led to low oxygen conditions in the oceans, and oxidation in ocean or atmosphere would have led to widespread dissolution of carbonates, thus shallowing of the CCD.

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The trigger for clathrate dissociation is not known; possible explanations include rapid warming of the intermediate ocean waters as a result of changing oceanic circulation patterns (e.g., Thomas, 1989, 1998; Kennett and Stott, 1991; Kaiho et al., 1996; Thomas et al., 2000; Dickens, 2001; Bice and Marotzke, 2002; Tripati and Elderfield, 2005; Nunes and Norris, 2006), continental slope failure as the result of increased current strength in the North Atlantic Ocean (Katz et al., 1999, 2001), sea-level lowering (Speijer and Wagner, 2002; Schmitz et al., 2004), the impact of a comet (Kent et al., 2003; Cramer and Kent, 2005) or other extraterrestrial body (Dolenec et al., 2000), explosive Caribbean volcanism (Bralower et al., 1997), North Atlantic basaltic volcanism (Eldholm and Thomas, 1993; Schmitz et al., 2004), or some combination of various of these possibilities. Arguments against the gas hydrate dissociation hypothesis as the only explanation for the CIE include low estimates (500– 3000 Gt C) for the size of the global oceanic gas hydrate reservoir in the recent oceans and thus even lower ones in the warm Paleocene oceans (e.g., Milkov, 2004; Cramer and Kent, 2005; Archer, 2007), and the magnitude and timing of the warming event (Cramer and Kent, 2005; Bowen et al., 2004). Moreover, the rise of the CCD by >2 km is much greater than estimated by assuming that 2000–2500 Gt carbon in methane was the sole source of carbon (Dickens et al., 1997). Finally, the full extent of the CIE and thus the exact amount of isotopically light carbon and its isotopic signature are still somewhat uncertain (discussion in Zachos et al., 2005). The many alternative hypotheses for the source of the isotopically light carbon include the body of a comet (Kent et al., 2003; Cramer and Kent, 2005), thermal liberation of methane from organic matter by igneous intrusions in the North Atlantic (Svensen et al., 2004) or from sediments in the Alaskan accretionary prism (Hudson and Magoon, 2002), burning of extensive peat deposits (Kurtz et al., 2003), oxidation of organic matter following desiccation of inland seas (Higgins and Schrag, 2004), and mantle-plume-induced lithospheric gas explosions (Phipps Morgan et al., 2004), possibly associated with the late Paleocene–early Eocene Canadian kimberlite province (Creaser et al., 2004). Several years ago it was suggested that the PETM might not have been a singular event, but only the most severe out of a series of global warming events coupled with carbon isotope anomalies and carbonate dissolution (called hyperthermals; Thomas and Zachos, 2000; Thomas et al. 2000). Hyperthermals have now been documented in upper Paleocene–lower Eocene sediment sequences in the southeastern Atlantic Ocean (ODP Leg 208, Shipboard Scientific Party, 2004) and the Pacific Ocean (ODP Leg 198, Shipboard Scientific Party, 2002), and in land sections in Italy (Galeotti et al., 2005; Agnini et al., 2005) and the United States (Lourens et al., 2005). Lourens et al. (2005) argued that the PETM occurred at a time of orbital modulation similar to that of an event ~2 m.y. later (Elmo) as well as at an event 1.2 m.y. after the Elmo event, called the X-event (Roehl et al., 2005). If the PETM was one of a series of events of varying magnitude, occurring at orbital periodicities

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(Lourens et al., 2005), its cause probably was not singular (e.g., a comet impact or a volcanic eruption), but intrinsic to Earth’s climate system. There is as yet no agreement on the linkage of the PETM to a specific orbital configuration (Cramer et al., 2003; Westerhold et al., 2007). It is outside the range of this paper to discuss climatic and biotic events during the PETM in detail (see e.g., papers in Aubry et al., 1998; Wing et al., 2003). Regardless of the exact cause(s) of the addition of isotopically light carbon to the environment, the question remains, What caused the extinction of deep-sea benthic foraminifera (see Thomas, 1998, for a review)? Shelf foraminifera were also affected by the end-Paleocene events (e.g., Speijer et al., 1997; Speijer and Schmitz, 1998; Alegret et al., 2005; Ernst et al., 2006), although neither shelf nor deep-ocean foraminifera were affected by the asteroid impact at the end of the Cretaceous (Culver, 2003). In contrast, calcareous planktic organisms did not suffer severe extinction at the end of the Paleocene, although planktic foraminifera (e.g., Kelly et al., 1996, 1998; Kelly, 2002) and calcareous nannoplankton show rapid evolutionary turnover and evolution of short-lived taxa (e.g., Bralower, 2002; Tremolada and Bralower, 2004). Major extinctions of planktic and benthic organisms in the oceans thus appear to have been decoupled (e.g., Thomas 1990b; Kaiho 1994a), in agreement with the above arguments that bentho-pelagic coupling was less close during the Late Cretaceous-Paleogene than it is today. What could have caused a global extinction in the deep sea? Possible causes include (1) low oxygenation, as mentioned by many authors (see review in Thomas, 1998), either as a result of increased deep-sea temperatures or as a result of oxidation of methane in the water column; (2) increased corrosivity of the waters for CaCO3 as a result of methane oxidation (Thomas, 1998; D. Thomas et al., 1999; Zachos et al., 2005) or invasion by CO2 from the atmosphere (Feely et al., 2004; Sabine et al., 2004; Orr et al., 2005); (3) increased or decreased productivity or expansion of the trophic resource continuum (e.g., Hallock, 1987; Hallock et al., 1991; Boersma et al., 1998; Thomas, 1998; Bak, 2005); or (4) a combination of several of these. In considering possible causes, it should be kept in mind that the extinction was global, affecting 30%–50% of species globally, and post-extinction faunas worldwide are low-diversity, dominated by small, thinwalled calcareous species or agglutinants (Thomas, 1998). Unfortunately, such assemblages might result from such differing environmental factors as high temperature, undersaturation with calcite, low dissolved oxygen levels, and either high or low food levels (Boltovskoy et al., 1991). In addition, such small individuals could be opportunistic taxa indicating a disturbed environment, as expected after a major extinction (e.g., Schröder et al., 1987). Low oxygen conditions have been well documented in marginal ocean basins such as the Tethys and northeastern periTethys, as shown by the occurrence of the extinction at the base of laminated, dark-brown to black sediments with high concentrations of organic matter (Gavrilov et al., 1997, 2003; Stupin and Muzylöv, 2001; Speijer and Schmitz, 1998; Speijer and Wagner, 2002; Alegret et al., 2005), as well as in New Zealand (Kaiho et al.,

1996). The record for open-ocean settings, however, is not so clear. At some pelagic locations, post-extinction benthic foraminiferal assemblages may be interpreted as indicative of a high food supply or low oxygenation because of the abundant occurrence of buliminid foraminifera (see Thomas, 1998 for a review; Thomas, 2003; Nomura and Takata, 2005). There is, however, no sedimentological or geochemical evidence for hypoxia or anoxia (e.g., high organic carbon content, laminated sediments). Organic carbon in the PETM clay layers may well have been oxidized post-depositionally (van Santvoort et al., 1996). Trace element information, including low levels of Mn in the PETM clay layer at ODP Site 926, suggests that low oxygen conditions might have been more widespread than is now accepted (Thomas and Röhl, 2002) but anoxia or even hypoxia was not a global phenomenon. The sedimentology at Maud Rise Sites 689 and 690 and Pacific Sites 1209 and 1210 clearly indicates persisting oxygenation (Thomas and Shackleton, 1996; Thomas, 1998; Kaiho et al., 2006), and Walvis Ridge sites show bioturbation through the Paleocene-Eocene clay layer (Shipboard Scientific Party, 2004). There thus should have been refugia available for deep-sea benthic foraminifera: many species are cosmopolitan and have early life stages that are easily dispersed by ocean currents (Alve and Goldstein, 2003). Even if locally the oceans were hypoxic or even anoxic, some regions should have remained suitable for benthic foraminifera, which have considerable tolerance for low oxygen levels (down to ~1 mg/L; e.g., Gooday, 2003). Return to more favorable conditions should have been followed by rapid re-establishment of foraminiferal populations (e.g., Hess and Kuhnt, 1996; Hess et al., 2001). Similarly, refugia should have existed for carbonate corrosivity, because at Maud Rise Sites 689 and 690 the carbonate percentage decreased, but not below ~65% (D. Thomas et al., 1999), and there is no clear clay layer (e.g., D. Thomas et al., 1999; Cramer et al., 2003). Dissolution along depth transects in the Pacific Ocean was not as severe as at Walvis Ridge (Colosimo et al., 2005; Nomura and Takata, 2005; Kaiho et al., 2006). Extinction levels at these sites are similar to those at sites where dissolution is intense (Zachos et al., 2005). If carbonate corrosivity and the rise of the CCD had been the main cause of the benthic foraminiferal extinction, organic-agglutinated foraminifera below the CCD would not have been affected, but these also suffered extinction (e.g., Kaminski et al., 1996; Bak, 2005; Galeotti et al., 2005; Kaminski and Gradstein, 2005). It seems improbable that decreasing productivity in the surface waters could have caused a major benthic extinction, given that such extinction did not occur at the end of the Cretaceous. Decreased productivity remains a possible cause (e.g., Kaminski and Gradstein, 2005), however, because productivity decrease at the end of the Cretaceous might have been short-lived, as argued above, but could have been more long-term during the PETM, with a duration of ~100 k.y. (e.g., Zachos et al., 2005). The effect of the PETM on oceanic productivity, however, was not consistent globally. Evidence from sections close to continental margins and in epicontinental basins indicates high productivity, leading to hypoxia or anoxia (Gavrilov et al., 2003; Speijer et al.,

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Figure 3. Faunal patterns during the benthic foraminiferal extinction (modified after Thomas, 1998). Numbers in the legend identify DSDP and ODP sites. Blue—Sites 689 and 690, Weddell Sea. Red—ODP Site 865, equatorial Pacific Black—ODP Site 926, western equatorial Atlantic Ocean. Green—Walvis Ridge, southeastern Atlantic Ocean including DSDP Sites 525 and 527 and ODP Sites 1262 and 1263. The net loss of diversity is similar at different sites outside the interval of dissolution, but the patterns of dominance differ by site. Species indicative of low-oxygen and/or low oxygenation (bi/triserial species) increased in relative abundance at Sites 865, 689, and 690 but not at the Walvis Ridge sites; these species were absent from Site 929. Nuttallides truempyi, a possible low-food indicator (Thomas, 1998) increased in abundance at Site 929 (after the interval of dissolution) and at Walvis Ridge sites. Abyssaminid taxa, which are small, thin-walled species that might indicate oligotrophy or might be opportunistic species, increased in abundance at the Walvis Ridge sites, especially the deepest sites, and at Site 929, and decreased at Site 865.

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abyssaminid species, % (low food? opportunists?) N. truempyi, % (low food) bi/triserial species, % (infaunal morphotype- high food)

1997; Speijer and Schmitz, 1998; Speijer and Wagner, 2002). Overall, however, it appears that productivity decreased in much of the open ocean (e.g., Kaminski and Gradstein, 2005), although proxy data on productivity are sometimes in conflict. For example, Bralower (2002) argued for oligotrophic conditions during the PETM at one location (Site 690 in the Weddell Sea), whereas Thomas and Shackleton (1996), Bains et al. (1999), Stoll and Bains (2003) argued for eutrophic conditions at that same location (but see also Thomas, 2003). At tropical Pacific Sites 865 and 1209 planktic foraminifera and calcareous nannofossils suggest oligotrophy, benthic foraminiferal accumulation rates, and faunal composition eutrophy (Kelly et al., 1996; Kelly et al., 2005; Thomas et al., 2000; Kaiho et al., 2006). At other locations (e.g., Walvis Ridge Sites 525, 527, 1262, 1263; Ceara Rise Site 929), benthic foraminiferal evidence indicates a decrease in productivity (Thomas, 1998), but this apparent decrease might have resulted from rising temperatures thus higher metabolic rates at a stable food supply (Thomas and Shackleton, 1996; Boersma et al., 1998; Thomas and Röhl, 2002). These different observations may suggest that the trophic resource continuum expanded, i.e., gyral regions became more oligotrophic, continental margins more eutrophic (e.g., Boersma et al., 1998). The overall evidence indicates neither global increase nor decrease in productivity (Fig. 3), but the benthic extinction was global, suggesting that productivity changes by themselves probably were not the main cause of the benthic foraminiferal extinction. A global feature of the PETM is warming, for which there were no refugia as far as we know: at all investigated sites, at all latitudes, in all oceans, there is evidence for rapid warming of benthic environments. This suggests that warming may have been the main cause of the global benthic foraminiferal extinction. Understanding whether this was indeed the case is important for predicting behavior of benthic global biota during possible future anthropogenic global warming, especially because there is evidence that the oceans are warming (e.g., Levitus et al., 2000). It is questionable whether benthic foraminiferal assemblages that today live in deep waters at fairly high temperatures, e.g., those of

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the Red Sea, could assist in interpretation of Paleocene-Eocene events. These assemblages live at high but constant temperatures (Thiel et al., 1987; Edelman-Furstenberg et al., 2001), and in a special setting (small ocean basin surrounded by desert). Unfortunately, the mechanism by which global warming would have caused the benthic extinction is not clear: the effects of present global warming on oceanic ecosystems and biogeochemistry are predicted to be significant, but they are not understood (e.g., Petchey et al., 1999; Laws et al., 2000; Archer et al., 2004; Sarmiento et al., 2004). Because temperature is such an important regulator of metabolic rates and affects food supplies of all components in the marine benthic ecosystem, a major and rapid temperature change would be expected to affect energy cycling within ecosystems, possibly affecting productivity by calcareous nannoplankton (Feely et al., 2004), predation by such foraminifer-specializing predators as gastropods and scaphopods (Hickman and Lipps, 1983; Culver and Lipps, 2003), and even the rate of evolutionary processes (Gillooly et al., 2004). Many deep-sea benthic foraminifera feed on organic matter and Bacteria and Archaea within the sediments, and temperature changes affect not only metabolic rates of prokaryotes, but also which species are most active, and which labile compounds are generated (Weston and Joye, 2005). Changing temperatures thus may have changed the compounds and the amount of labile organic matter available for foraminiferal feeding. How might we determine whether warming indeed caused the extinction? Highly detailed stable isotope and trace element records on benthic foraminifera, preferably along a depth transect, are not yet available because of the problems with analysis of biota across a major extinction, and the small size of post-extinction individuals. Such analyses, however, might establish whether warming occurred just before or at the beginning of the extinction. Investigation of the apertural configuration of benthic foraminiferal species groups could reveal changes in feeding strategy. Carbon isotope analyses of different benthic foraminiferal species may help elucidate the environmental preferences of now-extinct taxa (shallowdeep infaunal, epifaunal). Unfortunately, such high-resolution studies are not possible across the PETM at many locations, because of the severe carbonate dissolution and thus incompleteness of records (Zachos et al., 2005). High-resolution investigations of faunal and isotope patterns across PETM-like events in the early Eocene (Shipboard Scientific Party, 2004; Lourens et al., 2005; Roehl et al., 2005) may provide more detailed information on the possible linkage between global warming events and the extinction of deep-sea biota; because dissolution was less severe than during the PETM, high-resolution records may be obtainable. CONCLUSIONS In the present-day oceans, deep-sea benthic foraminiferal faunas are strongly influenced by primary producers at the sea surface, which constitute their food supply. During the early part of the Cenozoic, however, including the Cretaceous-Paleogene boundary and the Paleocene-Eocene boundary, extinctions in surface

and deep-ocean biota were decoupled. Benthic foraminifera may not have suffered severe extinction at the Cretaceous-Paleogene boundary because oceanic productivity as well as food transfer to the bottom recovered more quickly than previously thought, and/or because a larger fraction of food was produced by lithoautotrophic prokaryotes in the warm oceans of the Cretaceous and Paleogene. The major benthic extinction at the Paleocene-Eocene boundary is not easy to explain; possible causes include changing oceanic productivity, lowered oxygenation, and carbonate corrosivity, but none of these factors occurred globally, and survival in refugia followed by repopulation would have prevented extinction of cosmopolitan species. Global warming might have been the most important cause of extinction, but mechanisms are not understood. Bentho-pelagic coupling as we see it in today’s oceans may have originated by the Eocene-Oligocene transition, during an episode of growth of the Antarctic ice sheet. The gradual benthic foraminiferal turnovers during the Eocene-Oligocene, middle Miocene, and middle Pleistocene all occurred during episodes of cooling, and all included the loss of similar species that might have indicated a high or continuous food supply. Primary productivity, however, probably increased during these times, suggesting that transfer of food to the ocean floor was different during the warm Paleogene. Benthic faunal turnover during Cenozoic episodes of global cooling may reflect the increased seasonality of primary productivity and increased delivery of labile organic matter to the seafloor. ACKNOWLEDGMENTS Discussion in Karl Turekian’s coffee room at Yale helped in the formulation of the speculations presented in this paper. Research was in part funded by National Science Foundation grant EAR 0120727 to J.C. Zachos and E. Thomas. I thank Andy Gooday and Laia Alegret for there constructive reviews, which improved the manuscript. REFERENCES CITED Adams, J.B., Mann, M.E., and D’Hondt, S., 2004, The Cretaceous-Tertiary extinction: Modeling carbon flux and ecological response: Paleoceanography, v. 19, p. PA1002, doi: 10.1029/2002PA000849. Agnini, C., Fornaciari, E., Giusberti, L., Backman, J., Capraro, L., and Grandesso, P., Luciani, V., Muttoni, G., Rio, D., and Tateo, F., 2005, The early Paleogene of the Valbelluna (Venetian Southern Alps), in Field Trip Guidebook Leg 208 Post-Cruise Meeting: Padua, Italy: University of Padua, 33 p. Alegret, L., and Thomas, E., 2001, Upper Cretaceous and lower Paleogene benthic foraminifera from northeastern Mexico: Micropaleontology, v. 47, p. 269–316, doi: 10.2113/47.4.269. Alegret, L., and Thomas, E., 2004, Benthic foraminifera and environmental turnover across the Cretaceous/Paleogene boundary at Blake Nose, Western Atlantic (ODP Hole 1049C): Palaeogeography, Palaeoclimatology, Palaeoecology, v. 208, p. 59–83, doi: 10.1016/j.palaeo.2004.02.028. Alegret, L., and Thomas, E., 2005, Paleoenvironments across the Cretaceous/ Tertiary boundary in the central North Pacific (DSDP Site 465), the Gulf of Mexico and the Tethys: The benthic foraminiferal record: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 224, p. 53–82, doi: 10.1016/ j.palaeo.2005.03.031. Alegret, L., Molina, E., and Thomas, E., 2001, Benthic foraminifera at the Cretaceous/Tertiary boundary around the Gulf of Mexico: Geology, v. 29, p. 891–894, doi: 10.1130/0091-7613(2001)029<0891:BFATCT>2.0.CO;2.

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The Geological Society of America Special Paper 424 2007

A major Pliocene coccolithophore turnover: Change in morphological strategy in the photic zone Marie-Pierre Aubry Department of Geological Sciences, Rutgers University, 610 Taylor Road, Piscataway, New Jersey 08854-8066, USA, and Department of Geology and Geophysics, Woods Hole Oceanographic Institution, Woods Hole, Massachusetts 02543, USA

ABSTRACT The coccolithophores (or calcareous nannoplankton) have proven to be remarkably sensitive to changes in the earth system. However, their history is often expressed in terms of changes in species richness, a methodology that became suspect with the discovery of cryptic species through molecular techniques. To avoid this problem, I describe the extant coccolithophores in terms of morphostructural characters, tracing their changes through the Neogene. I conclude that these are regulated by a morphological strategy that favors small size of cells and coccoliths. I show that this strategy developed as a result of morphologic convergences in different lineages and in taxa inhabiting different strata of the photic zone. A long-term trend through the Neogene resulted in similar innovations in different lineages, and, also, in the loss of large and complex coccoliths. Superimposed on this trend, a short (~2 m.y.) Pliocene turnover involved both the loss of morphostructural groups that were successful through the Paleogene and Miocene, and a critical, permanent shift to smaller size in the dominant Family Noelaerhabdaceae. The life strategy exhibited by the extant calcareous nannoplankton is rooted in this turnover, so that the same morphological strategy (Pleistocene Morphological Strategy; PLMS) has regulated the coccolithophores since the latest Pliocene–earliest Pleistocene. It is possible that biologic pressure from the microplankton induced the shift to smaller cells; it is equally possible that the progressive change in size and fine structure of coccoliths through the Neogene is linked to an increase in the Mg/Ca ratio and decrease in Ca2+ concentration in what Stanley and Hardie have called Aragonite III sea. The Pliocene turnover was likely induced by glacial intensification in the middle Pliocene, and sustained by the progressive cooling from the warm early Pliocene to the cold Pleistocene. The PLMS thus results from the combined forcing of ocean chemistry and climatic change. The physiognomy of the extant coccolithophores, far from indicating their failure as a result of unfavorable seawater chemistry, demonstrates the remarkable adaptability of a group that evolved and first radiated in Aragonite II sea, thrived in Calcite II sea, and has “reinvented” itself in Aragonite III sea by adopting a unique morphological strategy. Keywords: Pliocene turnover, morphologic strategy, convergence, sea water chemistry.

Aubry, M.-P., 2007, A major Pliocene coccolithophore turnover: Change in morphological strategy in the photic zone, in Monechi, S., Coccioni, R., and Rampino, M.R., eds., Large Ecosystem Perturbations: Causes and Consequences: Geological Society of America Special Paper 424, p. 25–51, doi: 10.1130/2007.2424(02). For permission to copy, contact [email protected]. ©2007 The Geological Society of America. All rights reserved.

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INTRODUCTION The mass extinction events that punctuated the history of life have resulted in successive radiations among taxonomic groups of higher order, reflecting the sudden availability of new ecological opportunities (Sepkoski et al., 1981; Sepkoski, 1982a, 1982b; Raup and Sepkoski, 1982; Bambach et al., 2004, McGhee et al., 2004). Yet, as dramatic as any of these mass extinctions might have been, their long-term effect on subsequent radiations has been modulated by effective, though less obvious, turnovers. A case in point is the global turnover that occurred at the Paleocene-Eocene boundary (ca. 55 Ma), a mere 10 m.y. after the Cretaceous-Paleogene event (ca. 65 Ma). This turnover was the most significant evolutionary event of the Cenozoic, such that the evolutionary roots of many extant taxonomic groups extend back to it rather than to the dramatic CretaceousPaleogene boundary event (see Aubry et al., 1998). The main Phanerozoic mass extinction events were recognized from compilations of the stratigraphic ranges of genera and higher rank taxonomic entities among invertebrates (Sepkoski, 1982b). The amplitude of a mass extinction event such as the K-P event was, however, well known from the paleontological record of unicellular organisms (Bramlette, 1965), demonstrating that protists are as sensitive to environmental changes as are more complex organisms in the Kingdom Animalia. Indeed, studies of the calcareous plankton (planktonic foraminifera and coccolithophores) have been determinant in clarifying the significance of the extinctions at the Eocene-Oligocene boundary (Boersma and Premoli Silva, 1991; Hallock et al., 1991; Aubry, 1992), and in documenting the Paleocene-Eocene boundary turnover (e.g., Thomas, 1992; Kelly et al., 1996; Aubry, 1998; Steineck and Thomas, 1996; Kaiho et al., 1996; Ouda and Berggren, 2003). We thus can rely on the microfossils as indicators of significant events in earth history. The evolutionary history of the microplankton is commonly described in terms of changes in species diversity (e.g., Bown et al., 2004; Bown, 2005 for the coccolithophores). However, as for the planktonic foraminifera (de Vargas et al., 1999, Darling et al., 2001), the measurement of species diversity in the coccolithophores based on morphologic characters becomes increasingly controversial, as studies in molecular biology expand (Saez et al., 2003; Quinn et al., 2004). I concur with Young et al. (2005) that determination of species diversity through time is highly suspect, and I have explained (Aubry, 1989a, 1998) that, somewhat paradoxically, the history of generic diversity is more reliable than that of species diversity. Parallel with the recognition that the morphotypic approach is highly insufficient (see de Vargas et al., 2004; Saez and Lozano, 2005), other means of assessing diversity have been introduced, such as the morphostructure of coccoliths (Aubry, 1998) and the test size in the planktonic foraminifera (Schmidt et al., 2004) and diatoms (Finkel et al., 2005). I explore here a more comprehensive means of describing the diversity of the coccolithophores than in Aubry (1998). Using the idea that size and morphostructure of coccoliths and cocco-

spheres are combined in successive, well-defined patterns (morphological strategies) that are characteristic of successive ages, I describe a long-term trend toward decreasing size through the Neogene, and a rapid (2 m.y.) middle–late Pliocene turnover. The trend appears to correlate with changes in seawater chemistry. The Pliocene turnover, an hitherto unrecognized event that must have represented a major restructuring of the photic zone, likely results from progressive global cooling. The morphological strategy exhibited by the extant coccolithophores was completely established by the earliest Pleistocene. METHODOLOGY This study requires the comparison of extant and fossil material for two reasons. First, the living coccolithophores constitute, a priori, a model with which the fossil record can be compared and from which the biology and other characters of the extinct nannoplankton can be inferred (see for instance Young et al., 2005). Second, the living coccolithophores comprise a number of taxa that evolved only recently, during the Pliocene or later, and yet dominate extant communities at all latitudes and at most water depths. These taxa, of the Family Noelaerhabdaceae (which comprises the Pliocene and Pleistocene genera Emiliania, Crenalithus, Gephyrocapsa, Pseudoemiliania) are also dominant in upper Pliocene-Pleistocene assemblages. It is thus reasonable to assume that the morphologic traits of the extant coccolithophores can be extended back, at least through the late Pliocene. Conceptual Approach to Diversity Analysis in This Study Fossil coccolithophores belong to the Subclass Calcihaptophycidae (De Vargas et al., 2007), a group of essentially marine Haptophyceae (Edvardsen et al., 2000) that secrete organic and calcified scales forming a protective exoskeleton around the (commonly) spherical cell. The calcified scales and the mineralized exoskeleton are called coccoliths and coccosphere, respectively. Two types of coccoliths are produced. Holococcoliths are extracellular secretions, simply formed of minute, closely packed rhombs of calcite. Heterococcoliths are intracellular secretions that are expelled out of the plasmolemma. They differ from the holococcoliths in their elaborate construction of imbricated cycles of strongly modified, generally flattened, rhombs (Black, 1963; = elements). Each cycle consists of regularly arranged, morphologically (size and shape) and crystallographically identical elements. Cenozoic heterococcoliths are easily classified into a finite number of characteristic morphostructural groups. This morphostructural classification in turn supports a rigorous, objective generic taxonomy (Aubry, 1988a, 1989a, 1998). (The morphostructural groups among coccoliths are comparable, in a very broad sense, to the body plans in the Kingdom Animalia.) It would appear that holococcoliths and heterococcoliths represent, respectively, the motile and non-motile stages in the life cycle of the coccolithophores (Cros et al., 2000; Geisen et al., 2002).

Major Pliocene coccolithophore turnover Coccolithophore species in different extant and fossil genera (and higher taxonomic ranks) differ by several morphologic characters. The five most significant are (1) size and (2) morphostructure of (hetero)coccoliths, and (3) size of coccosphere and (4) number and (5) arrangement of constituent coccoliths (Fig. 1, Table 1). A survey of taxonomic and biostratigraphic syntheses (Tappan, 1980; Perch-Nielsen, 1985; Aubry, 1998; Young et al., 2003) shows that the broad Cenozoic evolutionary history of the coccolithophores can be described in terms of successive, characteristic combinations of these morphological characters. In other words, morphologic characters among the coccolithophores are not distributed at random, but organized in patterns that are time-specific and independent of phylogeny. I thus propose that each time-characteristic combination represents a morphological strategy, itself probably part of a global adaptive strategy, or biotic response to temporary combinations of forcing parameters. The history of morphological strategies in the coccolithophores is discussed elsewhere. The objective of this paper is to characterize the present-day morphological strategy, determine when it became established and discuss what forcing mechanism(s) may have caused it. It may then be possible to infer from the present-day strategy the forcing mechanisms that led to the establishment of earlier strategies. Procedures for Morphologic Analysis Coccoliths The morphostructure of coccoliths in different genera is described and discussed in the Handbook of Cenozoic Calcareous Nannoplankton (Aubry, 1984, 1988b, 1989b, 1990, 1999a, 2007a–d), which is the source of the morphostructural information used here. The measurements of coccoliths were taken from the original descriptions of taxa complemented, for extant forms, by additional measurements given in the taxonomic revisions by Kleijne (1991, 1992), Cros and Fortuño (2002), and the illustrated compilation of Young et al. (2003). Because coccoliths are threedimensional objects with very different shapes, it was necessary to establish a protocol for homogenous compilation of the data. The size data presented here concern the maximum diameter of the base of coccoliths. This is appropriate because this is also the maximum dimension in the majority of the Cenozoic morphostructural groups. However, some coccoliths (e.g., rhabdoliths, lopadoliths) are taller than wide. For extinct groups (e.g., sphenoliths) the concave side is regarded as proximal. For groups such as the (intracellular) ceratoliths, the cross section was measured rather than the total length. Extant Coccospheres Measurements regarding the size of coccospheres and their coccolith counts (i.e., the number of coccoliths forming a single coccosphere) are relatively few (except for Emiliania huxleyi), because of both the limited number of studies on extant coccolithophores and the low number of coccospheres per species reported in these studies. Comparisons of reports from different areas

27

and of illustrations by different authors indicate a considerable variability for some species, and a systematic quantitative study of intraspecific variations in the Family Rhabdosphaeraceae shows marked intraspecific variations (Kahn and Aubry, 2006). However, the amplitude of variation is consistently less than that between taxa. Size measurements and coccolith counts used here are from various sources, in particular Cros and Fortuño (2002), Kleijne (1991, 1992), Young et al. (2003), Winter and Siesser (1994), Yang and Wei (2003), and the original descriptions of many species. Tall coccoliths (such as rhabdoliths) form a complex coccosphere. The adjacent, proximal ends (or bases) of the coccoliths form an inner envelope or “inner coccosphere” whereas their distal ends delineate an outer envelope or “outer coccosphere” (see Kahn and Aubry, 2006). Complex coccospheres also occur in the case of pronounced dithecatism, as in Umbellosphaera species. The measurements given here are of the diameter of the outer coccosphere, and thus cannot be translated into cell diameters. Littoral and fresh-water taxa are not included in this study. Fossil Coccospheres I have searched the literature for published coccospheres in addition to the invaluable work of Mai (Mai, 1999; Mai et al., 1997, 1998, 2003) and entered a database recording the stratigraphic levels and localities from which the illustrated specimens were collected. All coccospheres considered here were well exposed. The total diameter of each coccosphere was measured and coccoliths were counted on the exposed side. The total number of coccoliths per coccosphere was estimated by doubling that count. This causes an error on the coccolith count, which increases with the number of coccoliths per coccosphere, but is always trivial for counts <50. When coccospheres are formed of a multitude of coccoliths, or when coccospheres are multilayered, the error increases dramatically. However, this study does not concern the individual taxon but groups of taxa. What is important here is whether coccospheres in these groups comprise fewer than ~30–40 coccoliths or more (see below). Temporal Constraints The extant calcareous nannoplankton comprise genera that have evolved at different times. Their temporal ranges are those determined from the fossil record, and keyed to the chronology of Berggren et al. (1995a) or when available to astrochronology (Lourens et al., 2005). Numerical ages by Gibbs et al. (2005) and Farrell et al. (1995) have been converted to the chronology of Lourens et al. The molecular ages on coccolithophores are limited at this time (Saez et al., 2003; Geisen et al., 2004) and still uncertain because of unverified calibration. Problematic are the ages of divergence of entire genera and families (e.g., Alisphaeraceae, Papposphaeraceae) that have no fossil record and for which molecular data are as yet unavailable. I have arbitrarily assigned a late Pleistocene age to these taxa, recognizing that some may be substantially older.

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Major Pliocene coccolithophore turnover MORPHOLOGICAL STRATEGY IN THE EXTANT COCCOLITHOPHORES Main Morphological Characteristics of the Extant Coccolithophores Although of apparent bewildering morphologic diversity, the extant coccolithophores are of strikingly general minuteness. The mean diameter of their coccoliths is 4 μm and the mean diameter of their coccospheres is 11.7 μm. The mean diameters of the coccosphere and coccolith of the overwhelmingly dominant cosmopolitan species Emiliania huxleyi are 6.9 μm and 3.5 μm, respectively. The mean diameter of the coccosphere of the deep-dwelling species Florisphaera profunda is 8.5 μm and the mean length of its coccolith is 3.15 μm (A. Kahn, personal commun., August 2006). In the most diversified extant family Syracosphaeraceae (37 formally described species), the mean diameter of the coccosphere is 11.8 μm and the mean diameter of the coccolith is 3 μm. The average size of all holococcoliths (66 “species” and morphotypes; produced by most species during the haploid stage, see above) is 1.7 μm and that of all holococcolith-bearing coccospheres is 9.71 μm.

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Nonetheless, the extant coccolithophores also include a few strikingly large species (see below under MT-Group 4). Collectively, the average diameter of their coccospheres is 20.75 μm, and that of their coccoliths is 10.1 μm. Individually, the average diameter of the coccosphere may reach 45 μm (as in Scyphosphaera apsteinii) and that of the coccoliths 15 μm (as in Coccolithus pelagicus). In other words, these taxa have coccospheres twice as large as those of the most dominant species (E. huxleyi and, in the deep photic zone, F. profunda) and the most diverse family (Syracosphaeraceae), and their coccoliths are three to five times larger than in the latter. The larger taxa also differ from the bulk of the smaller ones by other characters. One obvious difference is the coccolith count per coccosphere. The large species have a small number (average = 22) of coccoliths per coccosphere. In contrast, the average coccolith count in the Syracosphaeraceae (72) is more than three times larger. More subtle differences concern the structure of the heterococcoliths, in particular the arrangement of individual elements in adjacent cycles. In the large species, the elements are strongly overlapping whereas, as a rule, they are little or weakly imbricate in the smaller ones (Aubry, 1990, 2007a–d; see also illustrations in Cros and Fortuño, 2002; Young et al., 2003, 2005; and discussion below).

TABLE 1. CHARACTERS OF THE COCCOLITHS AND COCCOSPHERE OF INDIVIDUAL COCCOLITHOPHORE SPECIES USED TO DEFINE MORPHOLOGICAL STRATEGIES Coccoliths [1] Size: [1a] Amplitude of size range (= difference between the largest and smallest coccoliths at any given time) [1b] Mean size [1c] Mode [2] Morphostructure of discrete heterococcoliths (as shown in the base of a Paleogene rhabdolith): [2a] Number of cycles: generally varies from 1 to 6 [2b] Shape of elements in individual cycles: wedge-shaped (Cycle 3), trapezoidal, quadrangular (Cycle 1), rodshaped (Cycle 2), flat, thick, etc. [2c] Arrangement of elements in individual cycles. Elements may be non-imbricated (Cycle 1), disjunct (Cycle 2) or overlapping (imbricated; Cycles 3–5), spirally arranged (Cycle 6) [2d] Crystallographic orientation of elements (determined in light microscopy) [2e] Shape of cycles: circular, elliptical, lozenge-shaped spiral, etc. (see Fig. 7) [2f] Arrangement of cycles: concentrically arranged, spirally arranged, overlapping, etc. Coccospheres [3] Size: [3a] Amplitude of size range (= difference between the largest and smallest coccoliths at any given time) [3b] Average (mean) [3c] Mode [4] Number of coccoliths (= coccolith count) may vary considerably from a minimum of 4 to several hundreds [5] Structure of coccosphere: [5a] Monothecatism, dithecatism: coccospheres may be formed by one type or two types of coccoliths. Dithecatism occurs (as shown in Fig. 1 in Syracosphaera pulchra [MT-Group 1]) when the cell is protected by two concentric layers, each consisting of distinct coccoliths. [5b] Arrangement of coccoliths: the coccosphere may consist of overlapping (as shown in Oolithothus antillarum [MT-Group 2]) or non-imbricated (as shown in Papposphaera sp. [MT-Group 1]) coccoliths. [5c] Monomorphism, dimorphism, polymorphism: coccosphere may be uniformly made of similar coccoliths (e.g., Oolithotus antillarum) or may comprise coccoliths of different shapes either in equatorial position or at the flagellar pole (as shown in Fig. 1 in Calciosolenia murray [MT-Group 3] and Acanthoica sp. [MT-Group 2]). Note: See Figure 1. Characters that are readily available (bold italics), rarely available (italics), and unavailable (without italics) from the fossil record.

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Aubry

There are thus marked contrasts among extant species of the coccolithophores, contrasts that are largely concealed by the (superficially) considerable morphological diversity of the group. These contrasts have their roots in the time of origination of taxa. The extant coccolithophores comprise species of genera that have evolved at different times during the Cenozoic, as far back as early Paleocene (ca. 63 Ma; Coccolithus), or early Eocene (ca. 52 Ma; Pontosphaera, Scyphosphaera) and as recently as the latest Pleistocene (ca. 0.270 Ma; Emiliania). It also comprises Neogene genera that are rooted in Paleogene families (as in the Rhabdosphaeraceae). The analysis of the composition of the extant coccolithophores from a paleontologic perspective, based on the origination of their genera rather than in the usual terms of phylogenetic relationships, reveals that extant species of genera that originated during the same broad temporal interval share similar morphologic characteristics. Five Morphotemporal Groups in the Extant Coccolithophores The extant coccolithophores are readily described as comprising five natural morphotemporal (MT) groups. An MT-group is defined as a unique combination of (hetero)coccolith morphostructures ( = generic character), species richness, dominance, and time of generic origination (Figs. 1–3, Table 1). 1. MT-Group 1 constitutes the bulk of the extant calcareous nannoplankton and comprises well-diversified taxa about which the origin of the coccolith morphostructure remains generally obscure. It is exemplified by Syracosphaera and related genera (e.g., Ophiaster, Michaelsarsia; Syracosphaeraceae), Alisphaera, Pappomonas (Alisphaeraceae, Papposphaeraceae), and genera in many other incertae sedis families. In this highly diversified group, coccoliths are very small (<5 μm) and consist of a few cycles of thin, small elements that do not overlap or overlap little; coccospheres are generally, but not always, small (<12 μm), and all coccospheres—including the larger ones (>100 μm)—consist of a high to very high number of coccoliths (>40 to several hundreds) (Figs. 2 and 3). Coccospheres are often multilayered. Dithecatism and polymorphism are common (Fig. 1). Most taxa in this group have no known stratigraphic record, and molecular biology is much needed to determine their divergence time. The oldest stratigraphic record is Oligocene for the genus Syracosphaera, whose expansion is regarded as Pleistocene (Aubry 2007c); the divergence time of the species Syracosphaera pulchra is estimated as late Miocene (Geisen et al., 2004). 2. MT-Group 2 unites taxa that evolved during the Neogene and have diversified little, although several have achieved ecologic dominance. The morphostructures of their coccoliths are easily related to ancestral types. The group comprises species whose generally small (11 μm average) coccospheres consist of a large number (50 in average) of

small (4.8 μm average) coccoliths, most formed of a few elements. This group includes the 20+ species of the family Rhabdosphaeraceae, the two species of the Ceratolithaceae, and species of the genera Florisphaera (incertae sedis), Hayaster, Oolithotus (three dominant taxa in the deep water photic zone [Okada and McIntyre, 1977]), Umbellosphaera (dominant in the upper photic zone; Young et al., 2003), and Umbilicosphaera (Families Calcidiscaceae, Umbellosphaeraceae and Umbilicosphaeraceae, respectively) (Figs. 2 and 3). 3. MT-Group 3 consists of taxa with a very long stratigraphic record (Figs. 2 and 3). The coccosphere formed by 12 pentaliths in Braarudosphaera has remained unchanged since the Cretaceous. The long (up to 100 μm), fusiform coccospheres of Calciosolenia species (Family Calciosoleniaceae), which may contribute significantly to regional nannoplankton communities (e.g., Knappertsbusch, 1993), combine polymorphy with a high (>100) number of tiny (5 μm) scapholiths (Figs. 1[5c], 3B; scapholiths are tiny throughout the geological record). Except for their long stratigraphic record they would fit perfectly in MT-Group 1. 4. MT-Group 4 comprises species of genera that originated during the Paleogene and in which diversity has generally declined since the Pliocene (Fig. 2). Extant species of Coccolithus, Calcidiscus, Pontosphaera, Scyphosphaera, and Helicosphaera belong to this group. Most species have large (10–20 μm) coccospheres, consisting of few (12–40), large (up to 30μm) coccoliths (Fig. 3A and 3B) formed of several cycles of strongly overlapping elements. 5. MT-Group 5 includes taxa with tiny (4 μm) coccoliths and (6 μm) coccospheres but with a low (9–30) coccolith count (Fig. 3). The bulk of these taxa are the most recent (Pliocene and Pleistocene) representatives (genera Emiliania, Gephyrocapsa, and Crenalithus) of the family Noelaerhabdaceae (Fig. 2). In comparing the size of coccoliths and coccospheres, we note that: 1. MT-Groups 1–3 share strikingly similar features: all coccoliths are tiny (<5 μm), and most coccospheres are tiny (<12 μm) and consist of many coccoliths (>40 to hundreds). 2. MT-Group 4 contrasts sharply with MT-Groups 1–3. Most taxa have large coccospheres consisting of a small number (8–30) of generally large (8–20 μm) coccoliths. 3. MT-Group 5 occupies an enigmatic, albeit intermediate, position. Its coccoliths and coccospheres are generally small (as in MT-Groups 1, 2 and 3) but the number of coccoliths per coccosphere is low (8–26), as in MT-Group 4. Although extant coccolithophores include species (such as Syracosphaera pulchra [MT-Group 1], Oolithotus fragilis [MT-Group 2], and Emiliania huxleyi [MT-Group 5] that belong to distantly related molecular phylogenies (Saez et al., 2003, 2004) and, as importantly, occupy different ecological habitats in

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Figure 2. Temporal span of the Neogene coccolithophore genera. As genera are defined by morphostructure, this also represents the Neogene macroevolutionary history of coccolithophores. Thin bars for extinct genera, with thinnest bars for selected species. Thick bars for genera that are still represented. Numbers in parentheses indicate number of species in genus. Time scale of Berggren et al. (1995). Extant taxa without known fossil record are arbitrarily shown as having evolved ca. 0.200 Ma. Dashed line—range poorly established; *—informal genera (Aubry, 2007).

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Mean diameter of coccosphere Figure 3. Relation between the size of coccospheres and coccolith counts in the heterococcolith-stage of the extant coccolithophores (neritic species omitted). (A) Groups as described in text. (B) Same data but shown for specific taxa in each group. Note for instance that the coccolith count (mean = 50) per coccosphere for the extant rhabdolith-bearing species (MT-Group 2) falls within the average number of coccoliths per coccosphere for MT-Group 1. The genus Ceratolithus (Family Ceratolithaceae) (also MT-Group 2), which evolved in the earliest Pliocene from the relatively successful Amaurolithus, is another case in point. This taxon has a complex life cycle and produces numerous (>100), thin, ringshaped coccoliths in addition to a single massive horseshoe-shaped lith.

Major Pliocene coccolithophore turnover the photic zone, most are remarkably uniform in terms of size of coccoliths and coccospheres. How distinctive is this morphological uniformity? The answer resides in a comparison of the coccoliths and coccospheres of extinct and extant species in genera with a fossil record. MT-Group 1 is regarded as the norm because most extant species belongs to this highly diversified group. As demonstrated below, the morphologic uniformity exhibited by the extant coccolithophores reflects a distinctive morphological strategy as a result of widespread convergence in unrelated lineages. A Present-Day Morphological Strategy: Evidence from Coccoliths Homoplasy I discuss here three examples of morphologic convergence between coccoliths in MT-Groups 2 (rhabdoliths) and 4 (helicoliths) with coccoliths in MT-Group 1. Rhabdoliths. Rhabdoliths are coccoliths with a morphostructure typical of the family Rhabdosphaeraceae. The 20+ extant rhabdolith-secreting species are assigned to MT-Group 2. Rhabdoliths may have a distal elongated stem (styliform rhabdoliths s.l.) or a cupuliform distal process (sacculiform rhabdoliths). The Rhabdosphaeraceae evolved in the earliest Eocene (ca. 54.9 Ma) and was well diversified during the Paleogene, as it is today. However, its Neogene record is scarce. A decrease in size of both styliform and sacculiform rhabdoliths occurred from the Paleogene to the Neogene (Fig. 4) at the same time as their structure became simplified (Fig. 5A). Additionally the size of the styliform rhabdoliths has decreased through the Neogene (Fig. 5A), such that their size in extant species compares well with the size of coccoliths in MT-Group 1. In Palusphaera vandelii, the average diameter of the base and the average length of the stem are 1.5 μm and 5.7 μm, respectively (A. Kahn, personal commun., August 2006), similar to sizes in another extant species, Rhabdosphaera xyphos. There is no evidence (i.e., no fossil record, no molecular data) as yet that these two species evolved recently; however, by comparison with the fossil record of the group, the small size of these coccoliths suggest morphologic convergence with coccoliths of MT-Group 1. Helicoliths. The four extant species (Fig. 5B) of the family Helicosphaeraceae belong to MT-Group 4. The oldest fossil records of Helicosphaera carteri, H. wallichi, and H. pavimentum are, respectively, upper Oligocene (Aubry and Villa, 1996), upper Miocene (Young, 1998), and Pleistocene (Gibbs et al., 2005). Helicosphaera hyalina is known only from the living plankton, but Saez et al. (2003) have inferred from molecular biology an age of 10 Ma for its divergence from H. carteri. In addition to common morphostructure (Aubry, 1988a), molecular biology and life cycles show that these four taxa are phylogenetically related (Saez et al., 2003; Geisen et al., 2004). The size of the helicoliths produced by extinct species has varied broadly, ranging from 4 to 19 μm (Fig. 6). The living species secrete helicoliths that fall in the lower size range (5–10 μm). Further, there is both a sequential decrease in size of helicoliths

33

as new species evolved during the Neogene and a progressive reduction of the extent of the “distal cover” (a typical distal structure consisting of concentrically arranged, strongly overlapping elements; Fig. 5B). Thus, helicoliths of the extant H. pavimentum are half the size (length: 3.5–6 μm) of those of the long-ranging H. carteri; the distal cover is restricted to their central area in contrast to H. carteri, in which it extends over the whole distal face of the helicolith (Fig. 5B). Extant species of Helicosphaera (MT-Group 4) that have evolved more recently thus secrete helicoliths (1) that are notably smaller than in most of their extinct counterparts and (2) whose size compares well with that of coccoliths in MT-Group 1. Placoliths. Two morphostructural groups—placoliths and discoasters—have dominated the calcareous nannoplankton for the greater part of the Cenozoic. Discoaster-secreting species became extinct in the late Pliocene but placolith-secreting species are still among the most productive today (Fig. 2). The presence of extant placolith-bearing species in three MT-groups (Groups 2, 4, and 5) is a measure of the diversity and success of the placolith morphostructure. In contrast, other morphostructural groups are restricted to a single MT-group. Placoliths in different MT-groups exhibit the same Neogene morphologic trends as rhabdoliths and helicoliths: (1) a decrease in size and (2) a structural simplification. Size. As new genera of placolith-bearing species have evolved during the Neogene they have produced smaller coccoliths. This is seen among the Calcidiscaceae in MT-Group 2 (Fig. 7), in the successive divergence of Hayaster perplexus (7.5 μm), Oolithotus antillarum (6.5 μm), and Umbellosphaera tenuis (5.5μm). This size reduction is striking in the genus Calcidiscus (MT-Group 4). Three species are living today, well characterized by their distinct life-cycles and genetics (Geisen et al., 2004, Quinn et al., 2004). Their relations are as follows: Calcidiscus leptoporus (placoliths, 6.5 μm; coccosphere, ~10 μm) diverged from the large C. quadriperforatus (placoliths, ~7.5 μm; coccosphere, ~15 μm) ca. 11.6 Ma, from which the small C. braarudii (coccoliths, ~4 μm; coccosphere, 5.5 μm) diverged ca. 0.3 Ma. Among the Noelaerhabdaceae (Fig. 7) the late Pleistocene Emiliania huxleyi secretes placoliths (3.25 μm) that are half the size of those in the Pliocene–Pleistocene genera Crenalithus/ Gephyrocapsa/Pseudoemiliania spp. Conversely, the larger placoliths became extinct as shown by the latest Oligocene– early Pliocene sequence of extinctions of Reticulofenestra bisecta (12 μm), R. abisecta (12 μm), Coccolithus miopelagicus (15.5 μm) and Reticulofenestra pseudoumbilicus (6 μm). Simplification of Morphostructure. Parallel to the negative trend in placolith size, a reduction in the surface area of individual cycles of elements (the basic building block of a heterococcolith; Figure 1[2], Table 1) occurred through the Neogene. This is well illustrated in the family Calcidiscaeae, in which placoliths are circular and with a small central opening (Fig. 7). In Calcidiscus, the two shields are large, with strongly overlapping elements, and there is a small central structure. In Hayaster and Oolithotus, the proximal shield is markedly smaller than the distal one, the shield

34

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Mean Size (µm) Figure 4. Comparison of the frequency distribution of the mean size of Paleogene and Neogene rhabdoliths. (A) Styliform rhabdoliths (diameter and length). (B) Sacculiform rhabdoliths (diameter). Data from Kleijne (1992), Aubry (1999a), and Young et al. (2003).

elements are non-imbricated, and there is no central structure. In the related genus Umbellosphaera, the placolith structure is considerably modified, with the proximal shield extremely reduced and the elements of the distal shield very thin. In the Noelaerhabdaceae, the Neogene trends are the same but the structure was affected differently (Figs. 7 and 8). The significant changes concern the fine structure of the shields with a progression through time, from overlapping elements in both shields (R. pseudoumbilicus), to elements that are joined throughout their length in both shields (Gephyrocapsa and Crenalithus), to elements that are joined throughout their length in the proxi-

mal shield but only at their outer edge in the distal shield (Pseudoemiliania), to elements that are joined at their outer edge in both shields (Emiliania). Homeomorphy Although the coccolith type most abundantly produced by the living coccolithophorids is the placolith (because of Emiliania huxleyi), the most diversified morphostructural group is the caneolith, produced in the Syracosphaeraceae, which includes five genera and more than 48 species (i.e., one-quarter of presentday diversity; see Young et al., 2003). Compared with placoliths,

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Figure 5. Comparison of the history of size and structure in two unrelated families: (A) Rhabdosphaeraceae (Order Syracosphaerales Hay 1977) (Young et al., 2003) and (B) Helicosphaeraceae (Order Crepidolithales). Note that the overall size decrease is accompanied by a simplification of the structure of coccoliths. (A) Palusphaera vandeli and Rhabdosphaera xyphos (not shown) secrete the smallest known styliform rhabdoliths. dv—Distal view. pv—Proximal view. cs—Cross section. Stemless rhabdolith shown for Rhabdosphaera. (B) Note the decreasing area occupied by the distal cover (concentrically arranged elements) in younger species. Helicosphaera wallichi is only slightly smaller than H. carteri, and has a distal cover smaller than the latter but larger than H. pavimentum and H. hyalina, thus exhibiting intermediate characters. Also, H. pavimentum produces helicoliths with an almost regular (rectangular) outline, in contrast with the other three species in which the wing accentuates the typical pseudospiral structure of helicoliths. As a result, the surface area of the helicoliths of H. pavimentum is notably smaller than in other small species. Other Neogene species (e.g., H. orientalis, H. walberdorfensis) produced small helicoliths but always with a clearly marked wing (Aubry, 1990). Coccoliths drawn to scale. Duration of epochs arbitrary. Data from Aubry (1990, 1999a).

36

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Mean Length (µm) Figure 6. Comparison of the frequency distribution of the mean size ( = length) of extant (white) and extinct (black) helicoliths (genus Helicosphaera) (data from Aubry, 1990, and Young et al., 2003).

caneoliths may be seen as light constructions using numerous small, often rod-shaped elements that imbricate little or not at all. However, the morphostructural changes that placoliths underwent during the late Neogene (i.e., decreased imbrication and, ultimately, T-shaped elements) constitute a convergence toward the caneolith fine structure (Fig. 9A and 9B). Indeed, the fine structure of the placolith of E. huxleyi (and Crenalithus species) shows greater similarity to the fine structure of caneoliths than to any of the successful Paleogene and early Neogene placolith types, and, in particular, to Coccolithus pelagicus. Whereas the placolith of C. pelagicus and Gephyrocapsa oceanica are directly comparable, that of E. huxleyi departs strongly from the quintessential placolith that C. pelagicus symbolizes (Fig. 8). Homeomorphy between Emiliania huxleyi and Syracosphaera species involves not only the alteration of the fine structure of the Emiliania-placolith to mimic the Syracosphaeracaneolith, but also the reverse. Young et al. (2003) justly characterized the body coccoliths in the Syracosphaera molischii group as being placolith-like. I propose that this constitutes another case of homeomorphy in which the morphostructure of a caneolith is modified so as to mimic a placolith (Fig. 9C). A Present-Day Morphological Strategy: Evidence from Coccospheres For comparison between fossil and extant coccospheres, I have eliminated from the database the Paleocene coccospheres, and retained the 200 Eocene through Pleistocene coccospheres. All are placolith-bearing because the imbrication of adjacent placoliths confers on the coccospheres firmness, enabling their fossilization. Older coccospheres were eliminated to avoid any possible bias resulting from the major turnover that the calcareous nannoplankton underwent at the Paleocene-Eocene boundary (Aubry, 1998; Bown et al., 2004).

Eocene to Recent Placolith-Bearing Coccospheres The Eocene–Holocene database reveals three important facts (Fig. 10): 1. There is no apparent qualitative bias toward larger size among fossil placolith-bearing coccospheres. A substantial number of coccospheres are <5 μ. This decreases considerably concerns about a fossilization bias with respect to size throughout the geological record (Young et al., 2005). In fact because the largest coccoliths are also the least common at any stratigraphic level, if there were a bias it would be more toward larger coccospheres than toward smaller ones, as indicated by the lonely 45 μ large coccosphere (of Chiasmolithus grandis; Deflandre and Fert, 1954). 2. All coccospheres consist of a small number of placoliths, regardless of their diameter. The bulk of the coccospheres comprise between 8 and 20 coccoliths. We note that the largest coccosphere (45 μm) yields as few coccoliths (14) as the smallest (<5 μm) do. Although this is a single point, it contrasts sharply with the largest coccospheres in extant MT-Groups 1 and 2, which always consist of a large number of coccoliths (Fig. 3). 3. All coccospheres consist of a small number of placoliths regardless of taxonomy. It is remarkable that in species as different in fine structure and age as, among others, Coccolithus pelagicus (Paleogene–present), Reticulofenestra bisecta (middle Eocene–late Oligocene), R. pseudoumbilicus (early Miocene–early Pliocene), Gephyrocapsa spp. (late Miocene–present), and Pseudoemiliania lacunosa (early Pliocene–early Pleistocene), the number of placoliths per coccosphere varies between 6 and 22, with a mean between 12 and 18. Equally remarkable is that this trait occurs in remotely related lineages: C. pelagicus belongs to the order Coccosphaerales Haeckel 1894,

PALEOGENE

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Calcidiscus

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Gephyrocapsa

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Emiliania

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R. pseudoumbilicus

3.7 Ma

5 µm

Figure 7. Comparison of the history of size and structure of placoliths in two phylogenetically unrelated families, Calcidiscaceae (Order Coccosphaerales) and Noelaerhabdaceae (Order Isochrysidales). Distal (above) and proximal (below) views shown for each taxon. The proximal (concave) side of a coccolith is facing the cell. The distal (convex) side is in contact with seawater. Calcidiscaceae: The genus Umbellosphaera is generally not included in the family but is tentatively shown here as closely related. Noelaerhabdaceae: In R. pseudoumbilicus both shields consist of numerous, strongly overlapping, essentially parallelogram-shaped elements. In Gephyrocapsa and Crenalithus (not shown), the shield elements are also parallelogram-shaped but are poorly or non-imbricated. The proximal shield in Pseudoemiliania resembles that in Gephyrocapsa but the distal shield consists of T-shaped elements. In Emiliania, both shields consist of T-shaped elements. (See also Fig. 8.) Coccoliths shown at scale. Durations of epochs are arbitrary.

38

Aubry

Pseudoemiliania lacunosa

Emiliania huxleyi

ds

ds ps

Reticulofenestra pseudoumbilicus

Gephyrocapsa oceanica

ds

ds

ds ps

ds ps

R. bisecta

Coccolithus pelagicus

ds

ds

ds

ds ps

ds ps

ds ps

Figure 8. Comparison among placoliths in the Noelaerarhabdaceae (detail of Fig. 7). The fine structure (arrangement of the elements) of the shields in Reticulofenestra and (most) Gephyrocapsa is comparable to that in Coccolithus pelagicus (Family Coccolithaceae, Order Coccosphaerales) but markedly different from that in Pseudoemiliania and Emiliania.

emend. Young and Bown 1997; the other four genera to the order Isochrysidales Pascher 1910, emend. Edvardsen and Eikrem (in Edvardsen et al. 2000). It would be legitimate to argue that the fossil record is meaningless because the coccospheres of C. pelagicus (MT-Group 4), and Gephyrocapsa spp. and Emiliania huxleyi (MT-Group 5) show the same features, whether fossil or extant. Three lines of evidence indicate that this is not the case. First, the overall contribution of placolith-secreting species to the composition of the

calcareous nannoplankton through time must be considered. Second, it must be determined whether coccospheres bearing a multitude of placoliths (i.e., >30) have occurred in the past. It might be that for mechanical or other reasons placolith-bearing coccospheres have a characteristic low coccolith count. Third, it is useful to examine whether placoliths-bearing species still living but without known fossil coccospheres possess features distinctive from fossil coccospheres in the Eocene–Pleistocene database or fall in the same pattern.

Major Pliocene coccolithophore turnover

A

39

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S. pulchra

G. oceanica

S. pulchra

C. parvulus

S. molischii type 1

E. huxleyi

C

E. huxleyi type R

S. molischii type 4

Figure 9. (A, B) Comparison of the coccolith structure in extant species of Noelaerhabdaceae (Gephyrocapsa oceanica; Emiliania huxleyi; Crenalithus parvulus; Order Isochrysidales) and of Syracosphaeraceae (Syracosphaera; Order Syracosphaerales). (C) Homeomorphy between the placoliths of Emiliania huxleyi, Crenalithus spp., and the caneoliths of the Syracosphaera molischii group. The placoliths of G. oceanica differ fundamentally from the caneoliths of S. pulchra (A). In contrast, the fine structure of E. huxleyi compares well with that of S. pulchra, both consisting of thin rod-shaped elements (B). A few caneoliths (e.g., in the S. molischii group) also show a fine structure that is more akin to the placolith of E. huxleyi than to the typical caneolith of S. pulchra. Homeomorphy occurs also in the Noelaerhabdaceae, as for instance between E. huxleyi and extant Crenalithus spp. (e.g., Young et al., 2003, Pl. 1, Fig. 12; Pl. 3, Fig. 4; Pl. 20, Figures 1, 4, 6; Pl. 22, Fig. 8).

Diversity among Placolith-Bearing Species through the Cenozoic The extant calcareous nannoplankton includes 18 placolithbearing species distributed among nine genera (i.e., average of two species per genus, often differing by subtle characters), with at least three genera (Crenalithus, Gephyrocapsa, and Emiliania) extremely closely related (Young et al., 2003). Although mostly abundant, these species belong to genera that are little diversified today. As noted above, today’s diversity is not found among placolith-secreting families, but among the caneolith-secreting family Syracosphaeraceae. In contrast, placolith-bearing species (together with discoaster-secreting species) have dominated much of the Cenozoic nannoplankton. With regard to the coccolithophores, the Cenozoic can be said to be the “age of placoliths.” Variations around the placolith theme have produced 37 (mostly

long-ranging) genera and more than 300 (morphologic) species (i.e., average of eight species per genus). In terms of diversity, the placolith-bearing taxa lost their preponderance sometime during the Cenozoic (see below). Paleocene Placolith-Bearing Coccospheres The Paleocene record of placolith-bearing coccospheres is instructive. Early Paleocene coccospheres vary greatly in size (6.5–18 μm) and consist of a single layer of numerous placoliths; the greater the age, the more numerous. Thus the earliest Paleocene species Centumgemina simplex and C. scissuratus (Mai and Aubry, 2004) produced coccospheres with 120–250 placoliths (Mai, 2001). Coccospheres with 40–100 placoliths characterize the species of the early Danian genera Futyana, Praeprinsius, and Prinsius. Coccospheres in the Paleocene

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Coccosphere diameter (µm) Figure 10. Number of coccoliths versus diameter of 200 Eocene through Pleistocene placolith-bearing coccospheres. 1, Coccolithus spp.—C. pelagicus. 2, Calcidiscus leptoporus. 3, C. pelagicus. 4, Pseudoemiliania lacunosa. 5, Gephyrocapsa spp. 6, Reticulofenestra spp. Only Eocene through Pleistocene coccospheres are considered because there are no older direct phylogenetic links between the extant nannoplankton.

genera Cruciplacolithus and Toweius, among others, consist of fewer (<30) placoliths. We conclude that placolith-bearing coccospheres do not typically have a low coccolith count; the placolith count has changed through time. Other Extant Placolith-Bearing Coccospheres The post-Paleocene fossil record of coccospheres discussed here includes mostly coccospheres in the families Coccolithaceae and Noelaerhabdaceae, both of which evolved during the Paleogene (early Paleocene and early Eocene, respectively). However, there are also (essentially) Neogene families with placolith-bearing coccospheres that have extant representatives. These include the Calcidiscaceae (genera Oolithotus, Hayaster), Umbilicosphaeraceae (genus Umbilicosphaera), and Umbellosphaeraceae (genus Umbellosphaera) for which scattered fossil coccospheres are known. Interestingly, the coccospheres in these genera possess significantly more coccoliths than the coccospheres of the two Paleogene families. The coccospheres of the three ecologically dominant genera—Umbellosphaera, Oolithotus, and Hayaster (MT-Group 2)—consist of 14–30, 23–75, and ~100 placoliths, respectively. Similarly, some species of Umbilicosphaera (U. sibogae and U. annulus) have coccospheres with more than 60 placoliths.

A Present-Day Morphological Strategy: Additional Evidence Two other morphologic characters, polymorphism/dithecatism and holococcolith size, are also indicative of a present-day morphological strategy. Polymorphism/Dithecatism Polymorphism is the differentiation of the morphology of coccoliths according to position on a coccosphere (Fig. 1[5b, 5c]). It is interpreted as providing a variety of benefits, from flotation to protection of sensitive areas of a cell (Young et al., 2005). Polymorphism occurs in all MT-Groups except Group 5, but is widespread only in MT-Group 1. In dithecatism (Fig. 1[5a]), coccoliths of different morphologies are arranged in two concentric layers or (exo- and endo-) thecas. In the genus Rhabdosphaera (Family Rhabdosphaeraceae) the endotheca is formed by the stemmed rhabdoliths, the exotheca by rhabdoliths without stems. Whether with or without stems all rhabdoliths share the same basic structure. This is interpreted by some authors (e.g., Kleijne, 1992) as dithecatism, by others (e.g., Young et al., 2005) as polymorphism. A more striking form of dithecatism is obligatory in

Major Pliocene coccolithophore turnover the Syracosphaeraceae (MT-Group 1). In it the two concentric thecas differ not only in the morphology of their coccoliths, but also in their structure. The cell thus appears to be surrounded by two different coccospheres (Fig. 1[5a]). Dithecatism is interpreted as a means of protection in the case of loss of coccoliths (Young et al., 2005). I propose that it represents another form of morphologic convergence. In most species of Syracosphaera, the endotheca consists of cuplike coccoliths with their concave side turned outward whereas the exotheca consists of cuplike coccoliths with their concave side turned inward. Although the fit between the coccoliths of the endotheca and exotheca is poor, one coccolith of the exotheca straddling several coccoliths of the endotheca, dithecatism results in a two-layered coccosphere, comparable to a placolith-bearing coccosphere. As stated above, the placolith has been the successful morphostructural group for most of the Cenozoic, implying that two superposed layers of calcite is a strong (perhaps physiological, perhaps ecological) advantage in the Coccolithophores. Dithecatism may represent a convergence toward the placolith structure. A clear demonstration of this is found in the Syracosphaera molischii group, in which the morphology of the endococcolith is convergent with the placolith morphology (see above). In this group, the endo- and exococcoliths produce the best possible analogue to a placolith-bearing coccosphere. The exotheca of Syracosphaera anthos is also remarkably similar to the coccosphere of the placolith secreting Oolithothus antillarum (compare Young et al., 2003, Pl. 16, Fig. 4 with Pl. 6, Fig. 10). Holococcoliths With the exception of the order Isochrysidales, living coccolithophorids secrete holococcoliths as part of their complex life cycle (Cros et al., 2000; Geisen et al., 2002; Young et al., 2003, 2005). This is interpreted as a group strategy toward changes in nutrient availability (Cros et al., 2000). Although of extreme morphologic diversity, holococcoliths are not amenable to the same evolutionary analysis as heterococcoliths because of their undifferentiated structure. Nevertheless, the lack of correlation between size of coccospheres and coccolith counts (Fig. 11) shows that the holococcolith stage follows the morphological strategy determined from the heterococcolith stage. Large or small, holococcolith-bearing coccospheres are formed of numerous (50 to several hundred), tiny (<0.5 μm to 3.6μm) coccoliths. Moreover, morphologic diversity among holococcoliths is remarkably high although size is consistently small (<5 μm), and smaller than in extinct Cenozoic species (Fig. 12). Extinct Paleogene holococcoliths are locally abundant in the stratigraphic record (e.g., Aubry, 1988b). Although only a few species are known, they suggest a remarkable past diversity. However, contrary to the modern forms, Paleogene holococcoliths show a broad size range, of 21μm (from 4μm to 25μm; Fig. 12), that parallels the broad size range (28 μm) seen among Paleogene heterococcoliths (1.5–30 μm). The minuteness of holococcoliths in today’s ocean is revealing of the present-day morphological strategy.

41

Discussion Because the four families of Rhabdosphaeraceae (MTGroup 2), Helicosphaeraceae (MT-Group 4), Calcidiscaceae (MT-groups 2, 4), and Noelaerhabdaceae (MT-Group 5) have very different ancestral phylogenetic histories (indeed they belong to four different orders), and because the species discussed inhabit different strata of the photic zone, it would be very difficult to interpret the parallel changes in their morphology (size and fine structure) through time other than as convergences toward the morphologic traits exhibited by MP-Group 1. This implies that a strong forcing on morphology has occurred through the Neogene. This in turn implies a present-day morphological strategy. The occurrence of homeomorphy between Emiliania huxleyi and some species of Syracosphaera further shows that the forcing on morphology is so complex as to involve not only convergence toward MT-Group 1 (which is convergence toward the most successful morphology in terms of diversity) but also convergence of some taxa in MT-Group 1 toward morphology in MT-Group 5 (which is convergence toward the most successful morphology in terms of species dominance). Homoplasy may be the most conclusive evidence of a global morphological strategy among the extant calcareous nannoplankton. Although the number of fossil coccospheres is low, it is sufficient to show a variety of patterns. These patterns differ with time (e.g., early Paleocene versus late Paleocene–Eocene) and taxonomic position (e.g., Coccolithaceae and Noelaerhabdaceae versus Calcidiscaceae). The “coccosphere pattern” seen in the extant representatives of the placolith-bearing families in MT-Group 2 (Calcidiscaceae [Oolithotus, Hayaster], Umbilicosphaeraceae, and Umbellosphaeraceae) contrasts with that in MT-Group 4 (Coccolithaceae) and Group 5 (Noelaerhabdaceae), in which the coccolith count is characteristically low. The coccosphere pattern in the former is similar to that in the (non-placolith-bearing) taxa in MT-Group 1. This is easily interpreted in terms of convergence of coccosphere patterns in the more recent families. In summary, 11 of the 14 distinct families (Jordan et al., 2004) of marine extant coccolithophorids are united by a common morphological strategy. The other three families (MT-Group 4) are remnants of a successful past. They are relicts of a bygone era. The interlineage evolutionary convergence through the Neogene and homoplasy among extant taxa provide compelling evidence that microevolution in the coccolithophores is regulated by a distinctive, collective, morphological strategy. I have discussed here the marine coccolithophorids. The surprising fact is that their morphological strategy extends to the living coastal and fresh-water coccolithophorids (see below), themselves composed of the tiniest coccospheres with numerous tiny coccoliths (see Young et al., 2003, Pl. 9; I recognize, though, that no fresh-water calcareous nannofossil taxa are known).

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Mean diameter of coccosphere Figure 11. Relation between size of coccosphere and coccolith counts in the holococcolith-stage. (A) Note that larger coccospheres comprise a higher number of coccoliths than do smaller coccospheres. (B) Detail; coccospheres >200 μm omitted.

A PLIOCENE TURNOVER AMONG THE COCCOLITHOPHORES I have demonstrated a Neogene morphological trend toward smaller and lighter coccoliths in MT-Groups 2 and 4 (Figs. 2, 5, 7), foreshadowing the typical morphology of extant coccolithophores. This could indicate a progressive establishment of

the present-day morphological strategy. However, large coccoliths exhibiting morphostructures that have become lost (e.g., discoaster, ortholith, sphenolith; Fig. 2) or that are still represented but considerably smaller and/or less diversified (as in Scyphosphaera and Coccolithus) than in the past, dominated coccolithophorid communities throughout the Miocene, implying that the Miocene nannoplankton were regulated by morphological

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10 11 12 13 14

Holococcolith size (µm) Figure 12. Frequency distribution of the mean size (length) in extant and extinct holococcoliths. Data from Kleijne (1991) and Aubry (1988b).

strategy(ies) different from today. Thus, whereas changes operating over a long time scale have unequivocally contributed to its establishment, the present-day morphological strategy is more recent than the Miocene. Its origin must be sought in the Pliocene or Pleistocene, and, in all likelihood, is closely related to the evolution of taxa in MT-Group 5 (i.e., in the Noelaerhabdaceae). Specialization toward smaller size among the Noelaerhabdaceae is a Pliocene phenomenon, with the genera Crenalithus, Gephyrocapsa, and Pseudoemiliania rising almost simultaneously (Figs. 2 and 7). The structural differences between these taxa and Reticulofenestra have been mostly overlooked (Aubry, 2007a), the extinction of R. pseudoumbilicus being essentially regarded as the main Neogene evolutionary event in the family. Indeed the Last Appearance Datum (LAD) of the last large (up to 11 μm) species in the Noelaerhabdaceae at 3.7 Ma marks a shift to smaller size in that family (Gibbs et al., 2005; see also Young, 1990). This, however, should not be seen as an isolated evolutionary event in the family, but as a step in the initiation of a new strategy. The innovation of the Pseudoemiliania-placolith (ca. 4 Ma) and the LAD of R. pseudoumbilicus are equally significant as part of the long-term trend toward smaller, lighter coccoliths. Likewise, the LAD of R. pseudoumbilicus was not an isolated extinction event among the coccolithophores, but the first in a 1.75-m.y.-long sequence of extinctions at a high rate of one species per 0.175 m.y. through the middle (3.6–2.59 Ma) and late

(2.59–1.81 Ma) Pliocene (Piacenzian–Gelasian Stages) (Fig. 2). It was accompanied by the loss of sphenoliths, a morphostructural group that evolved in the early Paleocene, became a common element of the Eocene through middle Miocene nannoplankton, and then dwindled in diversity and frequency (Aubry, 1989b). We note that the two species that lived until ca. 3.54 Ma were very small (3 μm in diameter and height for the smallest Sphenolithus neoabies, compared to 5.5 μ in diameter and 9.5 μ in height for the lower–middle Miocene [17.7–13.5 Ma] S. heteromorphus). The LAD of R. pseudoumbilicus was also followed by the disappearance of the discoasters, the morphostructural group that together with placoliths dominated coccolithophore communities for over 55 m.y. (late Paleocene to Pliocene). Their disappearance involved the successive LADs of Eudiscoaster variabilis (ca. 12–2.9 Ma), E. tamalis (3.97–2.80 Ma), E. surculus (ca. 9.3– 2.5 Ma), E. asymmetricus (ca. 2.45 Ma), E. pentaradiatus (ca. 9.3–2.39 Ma), and E. brouweri (ca. 9.3–1.95 Ma, shortly before the Pliocene-Pleistocene boundary). It is remarkable that these extinctions were ordered according to decreasing size, in accordance with the general trend described among discoasters by Stanley and Hardie (1998). The close association of the LADs of R. pseudoumbilicus and Sphenolithus spp. was recognized by Gibbs et al. (2005) as indicative of a critical mid-Pliocene event in the evolution of the coccolithophores. However, these two events should not

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be seen in isolation. Both the extinctions of Eudiscoaster spp. and of Sphenolithus spp. mark the end of lineages that had been among the most prolific, in terms of both individual abundance and diversification, at least from the Eocene through Miocene. Major evolutionary events affected the coccolithophores between ca. 4 Ma and 1.95 Ma. They were part of a ~2-m.y.-long turnover bracketed by rapid innovations in the Noelaerhabdaceae and the extinction of the discoasters. The wave of late Pliocene extinctions continued into the Pleistocene (e.g., the well-known LADs of Calcidiscus macintyrei, Helicosphaera sellii, Geminilithella rotula, and Pseudoemiliania lacunosa) accompanied by rapid species turnovers (e.g., in Gephyrocapsa, Samtleben, 1980) and morphologic innovations such as the recent (0.29 Ma) Emilianiaplacolith. The events that determined the Pleistocene evolution of the Coccolithophores belong to the Pliocene turnover, however: (1) 1.95 Ma marks the loss of one of the two prominent morphostructural groups of the Cenozoic, and (2) since ca. 4 Ma only lineages of tiny calcareous nannoplankton (e.g., Gephyrocapsa, Crenalithus, and Rhabdosphaeraceae) have radiated. The innovation of the placolith with T-shaped elements, clearly derived from the Reticulofenestra-placolith, would lead to the success of E. huxleyi through Gephyrocapsa (Figs. 7, 8). I conclude therefore that the present-day morphological strategy is rooted in the Pliocene turnover, and essentially characterizes the Pleistocene. I refer to it below as the Pleistocene morphological strategy (PLMS), which favors tiny cells and coccoliths. FORCING MECHANISM ON THE PLMS Parallel to the morphostructural changes in coccoliths, from which the timing and modalities of the PLMS can be determined, coccospheres also underwent a change in size and number of coccoliths, which cannot be recovered directly from the paleontologic record but can be inferred from long-lived taxa. Each phenomenon requires explanation, which can be sought among biotic and abiotic mechanisms alike. Biological Pressure The immediate questions that arise concerning the driving force(s) behind the establishment of the PLMS are difficult to address because the raison d’être of the coccoliths remains unknown. Young (1994) and Brownlee and Taylor (2004) have reviewed the role of coccoliths and calcification in the coccolithophorids, respectively. The latter authors concluded that basic questions are unanswered as yet. Aside from the physiological role of calcification, the ecological factors that favor minute cells also have yet to be determined. If small cells compete better for light and nutrients than larger cells (Riegman et al., 1993), the PLMS that favors small size in the coccolithophores may reflect the biological pressure imposed by the planktonic foraminifera and the diatoms. The planktonic foraminifera dwelling in the photic zone possess photosynthetic symbionts (Hemleben et al., 1988), which may place them in direct competition with cocco-

lithophorids for light. Schmidt et al. (2004) have shown that, in tropical and subtropical areas, the maximum test size has considerably increased since the late Miocene. Their Figure 1 (p. 208) suggests an even more pronounced size increase since the midPliocene. Diatoms became considerably more diverse since the late Miocene. Owing to the storage of triacylglycerides, planktonic diatoms may easily outcompete other planktonic groups (Lombardi and Wangersky, 1991; Sicko-Goad et al., 1988). Accepting a biotic pressure scenario, however, would imply that the strategy of the coccolithophorids is controlled by the strategies of other (larger) groups, without themselves influencing their counterparts. The mean size of diatoms decreased through the Cenozoic (Finkel et al., 2005), implying perhaps that strategies evolved similarly in different phytoplanktonic groups in response to physical and chemical changes in their shared environment. Seawater Chemistry The long-term trend and the Pliocene turnover in the coccolithophores occurred during Aragonite III (Stanley and Hardie, 1998; see also Stanley and Hardie, 1999). It is now well recognized that seawater chemistry has changed through time from the interplay between riverine input and growth of oceanic crust (Hardie, 1996; Lowenstein et al., 2003; Holland, 2005). Using a model predicting secular variations of the Mg/Ca ratio, Stanley and Hardie (1998) have distinguished five alternating episodes during which massive biomineralization was either aragonitic (Aragonite I, II, III) or calcitic (Calcite I, II). When the Mg/Ca ratio is low, the concentration in Ca2+ is high, and massive aragonitic biomineralization is inhibited. In contrast, when the Mg/Ca ratio is high, the concentration of Ca2+ is low, and aragonitic biomineralization is facilitated. Reef building and biocarbonate sedimentation have thus changed in these aragonitic and calcitic seas. Since the late Paleogene the seawater chemistry (i.e., high Mg/Ca ratio and low Ca+2 concentration) has favored organisms that build their skeletons out of magnesium calcite and aragonite (i.e., scleractinian corals and Halimeda). Stanley and Hardie (1998) chose the coccolithophores, whose coccoliths consist of low magnesium calcite (Lowenstam and Weiner, 1989; Stoll et al., 2001), as evidence of the role of seawater chemistry on biomineralization. They noted that highly diversified, robustly calcified coccolithophorids contributed to massive chalk deposition in Hardie’s predicted Calcite II sea, whereas in the subsequent (predicted) Aragonite III sea chalk deposition had ceased because the calcareous nannoplankton were considerably less calcified. They reached the conclusion that secular variations in the Mg/Ca ratio and Ca2+ concentration of seawater did control the evolutionary history of the calcareous nannoplankton (as of other taxa). This conclusion is supported by the observation that normal calcification in Emiliania huxleyi is rigorously dependent on the calcium and magnesium concentrations of seawater (Herfort et al., 2004). The convergent Neogene trends that I have described in several morphostructural groups have resulted through time in

Major Pliocene coccolithophore turnover a decreasing amount of calcite necessary to produce a coccolith. For comparison, the weight of a placolith of Emiliania huxleyi is ~2 pg and that of Gephyrocapsa oceanica varies between 10.88 and 54 pg (Beaufort and Heussner, 1999; Young and Ziveri, 2000; Baumann, 2004; Beaufort 2005), but that of the early–middle Miocene species Coccolithus miopelagicus (LAD at 10.97 Ma) is ~1000 pg. Thus, there is support for Stanley and Hardie’s conclusion (1998) that the Neogene history of the calcareous nannoplankton is explained by changing seawater chemistry, even if in the detail the history of changes in size through the Miocene is more complex than these authors have inferred. It is plausible that the Calcite II calcareous nannoplankton have adapted to the Aragonite III seawater chemistry by changing their morphologic strategy. The establishment of the new strategy has first been progressive, with the extinction of the species with the heaviest coccoliths, but more importantly, with the concomitant appearances of simplified morphostructures clearly derived from earlier, complex ones. This has been described above for the rhabdoliths, helicoliths, and placoliths (Figs. 5, 7, 8). Taxa would have evolved with the physiological adaptations necessary to adjust to the lower Ca2+ concentration and the rising Mg/Ca ratio. Progressively and collectively, they became part of the PLMS. Inevitably, not all morphostructures could adjust to the rising Mg content in seawater. The discoaster- and sphenolithsecreting genera thus became extinct. Sphenoliths consist of conical elements that radiate from a short central axis, a rather rigid morphostructure if there is one. There is also little structural flexibility in discoasters because of the characteristic radial arrangement of their arms. The brief Pliocene occurrence of Discoaster tamalis (3.97–2.80 Ma) can be seen as a failed attempt to conform to the small-size strategy. Except for the middle Miocene taxon Catinaster, departures from the basic discoaster morphostructure have resulted in short-lived taxa (e.g., D. araneus and D. anartios during the Paleocene-Eocene boundary thermal maximum [PETM]; Aubry, 1999b; Kahn and Aubry, 2004). In contrast, morphostructures such as placoliths and rhabdoliths are easily modified to produce a large array of adaptive variations, as possibly seen in the Neogene Aragonite III sea compared to the Paleogene Calcite II sea. The caneolith (the most successful morphostructure in the extant coccolithophores as measured by species richness) is also a remarkably flexible structure (leading to homeomorphy with the placolith of E. huxleyi). It is also possible that the Pliocene turnover was partly caused, and the initiation of the PLMS maintained and reinforced, by the sustained riverine input of Mg after 2.8 Ma, as a result of erosion linked to high-frequency glacial-interglacial cycles and to Himalayan tectonics (Raymo et al., 1988; Capo and De Paolo, 1990; Hodell et al., 1990), and indicated by an increase in the 87Sr/86Sr ratio ca. 2.8 Ma (Farrell et al., 1995; McArthur and Howarth, 2005). The biosynthesis of tiny aragonitic coccoliths by a few living coccolithophorids (Manton and Oates, 1980; Thomsen and Buck, 1998; Cros and Fortuño, 2002) may demonstrate the global influence of seawater chemistry on calcification by the nannoplankton. This interpretation must be considered with cau-

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tion, however, as Hart et al. (1965) reported on a (single) Upper Cretaceous aragonitic coccolith (in Calcite II sea) and because no systematic study of the mineralogic composition of coccoliths seems to have been undertaken. Other Factors The secular changes in the Mg/Ca ratio of seawater may explain the long-term shift and mid-Pliocene turnover toward smaller coccoliths. However, more significant than the size of the coccoliths in the establishment of the PLMS was the overall reduction in cell size, because it implied a major physiological adjustment (Ryther, 1969; Banse, 1982; Riegman et al., 1993). This is because a twofold decrease in diameter of a cell (generally a sphere) results in an eightfold decrease in volume. Thus, whereas the cell of Coccolithus pelagicus is 5000μ3, that of E. huxleyi is 60 μ3. The forcing mechanism on morphological strategy in the coccolithophores may therefore not be linked to seawater chemistry, or it may be linked to it but in combination with other factors. The Pliocene turnover was synchronous with major climatic events. Middle Pliocene (Gelasian Stage) events were determinant in the history of the earth system (Zachos et al., 2001; Pillans and Naish, 2004; Lourens et al., 2005). The extinction of Sphenolithus spp. and Reticulofenestra pseudoumbilicus (ca. 3.7 Ma) coincided with glacial intensification between 3.8 and 3.6 Ma, and may have resulted from it (Gibbs et al., 2005). The cooling that induced the final stage of Northern Hemisphere glaciation at 3 Ma, the shift at 2.8 Ma, and the high climate variability during the late Pliocene and early Pleistocene had dramatic effect on continental sedimentation, landscapes, and terrestrial vegetations around the world. It is possible that high climatic variability caused sudden changes in oceanic biological regimes as well, possibly resulting in the rapid succession of global acmes and replacements of taxa between 3.7 and 2.5 Ma and since (Gibbs et al., 2005 and references therein). While elevating the concentration of Mg in seawater, increased transient runoff and erosion (see above) would also have increased the input of nutrients into the ocean, contributing to the radiation of small taxa capable of rapid growth such as Gephyrocapsa spp. (G. oceanica causes massive blooms in today’s ocean [Blackburn and Cresswell, 1993]). The genus Gephyrocapsa evolved in the late Miocene (Pujos, 1987; Triantaphyllou, 2000) but did not radiate until the mid-Pliocene. This geological change combined with a gradual cooling in the transition from the warm early Pliocene to the cold Pleistocene (Ravelo et al., 2004) may explain the Pliocene turnover and extinction that followed. Sustained climatic variability during the Pleistocene and continued input of nutrients to the ocean may have contributed eventually to the evolution of the ultimate opportunistic species Emiliania huxleyi, known for its very extensive blooms (Balch et al., 1991; Holligan et al., 1993). In this scenario, the ecological success of MT-Group 5 is easily explained. Its taxa may have adapted to rapid fluctuations in nutrient conditions by decreasing cell size (convergence with

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MT-Group 1) but retaining the low coccolith count that characterized the ancestral Reticulofenestra (as taxa in MT-Group 4). The long-ranging Braarudosphaera bigelowii (MT-Group 3) also fits in the LPMS because of its equal capacity to grow rapidly under high-nutrient conditions, thus also producing massive blooms (as first identified by Gran and Braarud, 1935). There are, however, other possible explanations for the success of MT-Group 5 and, in particular, of E. huxleyi. Species in MTGroup 5 are characterized by the biochemical production of alkenones. This group of lipids has a significant, albeit poorly understood physiological role in Emiliania huxleyi (Prahl et al., 2003). Alkenone biosynthesis might confer a strong physiological advantage to MT-Group 5 species, thus compensating for their incomplete adherence (low coccolith count) to the typical present-day morphological strategy. Further, Emiliania huxleyi differs from most other species, including its stem group, in continuously secreting and shedding coccoliths or in making a multi-layered coccosphere (Paasche, 1999). This feature, also known in another placolith-bearing species, Cruciplacolithus neohelis (MT-Group 4) (Fresnel, 1986), can be seen as another way of producing a coccosphere with a multitude of coccoliths, as in MT-Group 1, and constitutes further evidence of convergence among phylogenetically unrelated taxa. DISCUSSION This is the first time that the coccolithophores are described in terms of a morphological strategy, although successive Paleogene morphological strategies were unknowingly described by Aubry (1998). This new evolutionary interpretation, not in terms of species richness (Bown et al., 2004) or linear trajectories (Katz et al., 2004), but in terms of a succession of well-characterized morphological strategies, will require a reexamination of the mechanisms that have driven the evolutionary history of the coccolithophores. This study shows that abiotic environmental factors can play a very strong selection pressure on phylogenetic evolution. Several abiotic factors, whose forcings may be initiated at different times but that ultimately operate in concert, impose the progressive establishment of a strategy. Clearly the PLMS reflects a mosaic of convergences for small size initiated at different times and for different reasons, and points to the overall pressure that abiotic forcing imposes on biologic evolution in the photic zone rather than illustrating distinctive solutions to biotic pressure (see Worm et al., 2002; and in agreement with Schmidt et al., 2004). If morphological strategies are direct expressions of life strategies of extinct organisms, identifying successive morphological (life) strategies will shed new light on the evolution of the earth system. The problem is that inference about the evolution of the earth system can probably only be made within the framework of the strategy characteristic of the time of interest. Smaller and larger cells have different physiology and biology (Raymont, 1980). In this perspective, it is doubtful that the model system approach based on the extant taxon E. huxleyi (Westbroek et al.,

1993) can be applied beyond the earliest Pleistocene, not even for the early Pliocene ocean, let alone the Eocene ocean. An example of experimental evidence contradicted by “old” fossil evidence is the effect of CO2 concentration on calcification. Extant coccolithophorids reduce calcification in experimental increases of CO2 levels (Riebesell et al., 2000). However, Paleogene coccolithophorids produced large, robust coccoliths under pCO2 levels much higher (Pearson and Palmer, 2000) than in today’s experiments. The largest and heaviest Cenozoic coccolith (Chiasmolithus gigas) is up to 30 μm in diameter. With an estimated weight of 1600 pg, it is three times heavier than the heaviest coccoliths secreted today (Scyphosphaera apsteinii and Pontosphaera discopora, with an estimated weight of 540 pg; Triantaphyllou et al., 2004), over a hundred times heavier than Gephyrocapsa oceanica (weight estimates of 10.88–16.77 pg; Baumann, 2004), and 800 times heavier than Emiliania huxleyi and Florisphaera profunda (2 pg; Young and Ziveri, 2000; Baumann, 2004; Beaufort, 2005). Emiliania huxleyi is unlikely to serve as a reliable model outside of the PLMS because its physiological requirements evolved under this strategy. Although their physiologic requirements must have changed through time, species like Coccolithus pelagicus (MTGroup 4) are more likely to offer insights on past strategies and biological and physiological requirements. Paleobiological investigations need to address questions relevant not only to the size of coccoliths, but also to their finestructure, and, perhaps most importantly, relevant to cell volume and coccosphere weight. The different growth response to a changing Sr/Ca ratio of large cells (C. pelagicus, H. carteri, C. leptoporus [MT-Group 4] compared to small cells (e.g., G. oceanica, E. huxleyi [MT-Group 5], S. pulchra [MT-Group 1], A. robusta [MT-Group 2]) seen in culture experiments (Stoll and Ziveri, 2004; see also Ziveri et al., 2003) may simply reflect the different strategies between the two groups, the small species belonging to the PLMS, the larger one a remnant of past strategies. These experiments support the interpretation I offer of the morphological diversity and species composition of the extant coccolithophores. The taxa of MT-Group 4 are the few extant representatives of genera that contributed most to Paleogene through Miocene diversity. Their coccoliths represent basic morphostructures that supported Paleogene and/or Neogene (pre-Pliocene) radiations (Perch-Nielsen, 1985; Siesser, 1998; Aubry, 1990). Coccolithus pelagicus, Pontosphaera spp., and Scyphosphaera apsteinii are thus best interpreted as relict species representing past strategies. Helicosphaera carteri is also a relict species, but morphologic characters in the genus have changed through the Neogene in a convergent manner so that the (presumably) more recent species fit closely the present-day morphological strategy. This is true also of Calcidiscus. The young species C. braarudii fits the present-day morphology as beautifully as H. pavimentum does. Coccolithus pelagicus is a long-ranging taxon that was broadly distributed throughout the Paleogene and until recently (McIntyre and Bé, 1967; Haq and Lohmann, 1976; Geisen et al., 2004), implying a remarkable physiological adaptability (or neu-

Major Pliocene coccolithophore turnover trality). It is currently represented by three subspecies (Geisen et al., 2002; Parente et al., 2004), one restricted to Arctic waters (C. pelagicus ssp. pelagicus, coccolith length: 6–10 μ), the other two (C. pelagicus ssp. braarudii, 9–15 μ, and C. pelagicus ssp. azorinus, ~14 μ) occurring in temperate water. The divergence between subspecies pelagicus and braarudii is dated at 2.15 ± 0.57 Ma (Saez et al., 2003), suggesting that C. pelagicus escapes the selective pressure imposed on other taxa. Neocruciplacolithus neohelis also belongs to MT-Group 4, and it may be significant that in culture experiments it does not incorporate Mg in its coccolith when the Mg/Ca ratio of seawater rises whereas Pleurochrysis carterae and Ochrosphaera neapolitana do (Stanley et al., 2005). The latter two species are nearshore taxa, well adapted to changes in Mg concentration of seawater. I have noted above that littoral species follow the same morphological strategy as oceanic species. In light of the Stanley et al.’s observations, it may be that the Pliocene turnover resulted in a morphological strategy of oceanic species converging toward that of littoral species. For thermodynamic reasons it may be less energy costly to secrete smaller coccoliths when the Mg/Ca ratio of seawater is high. The long-term size reduction of coccoliths has been interpreted as symptomatic of a declining group. It is thus relevant to ask whether the mid-Pliocene turnover has contributed to the demise of the coccolithophorids. The living coccolithophorids are known for their massive blooms (see discussions in Tyrrell and Merico, 2004; and Balch, 2004) during which enormous amounts of calcium carbonate are produced. Thus, Stanley and Hardie’s description (1998; 1999, p. 5) of “evolutionary osteoporosis” of late Neogene coccolithophorids is startling. These authors contend that the rise of the Mg/Ca ratio and concomitant decrease in Ca2+ concentration led to their failure, measured by (a) their inability to contribute to massive chalk formation and (b) reduced calcification as seen in the general decrease in size of coccoliths, and the thinly encrusted cells. Basically, the calcareous nannoplankton changed from being a hypercalcifying taxon in Calcite II sea to a fragile, undercalcifying and undercalcified taxon in Aragonite III sea. This alleged failure could be ultimately seen as a cause, among others, of the domain shift hypothesis (Karl et al., 2001), which suggests that community structure in the phytoplankton may experience a global shift, with chlorophyll b (prokaryotes) primary producers replacing chlorophyll a primary producers (coccolithophores produce chlorophyll a and c). It is correct that, at first glance, the late Neogene coccolithophorids compare poorly with their older relatives. Nonetheless, I would argue that, contrary to superficial evidence (size), the calcareous nannoplankton did not fail in Aragonite III sea. Coccoliths of Aragonite III sea have as symmetrical and elaborate morphostructures as coccoliths of Calcite II sea (Cretaceous and early Paleogene) and coccospheres of Aragonite III sea protect cells as efficiently as coccospheres of Calcite II sea did. On the other hand, increasingly reported malformations of coccoliths and coccospheres (Kleijne, 1990; Giraudeau et al., 1993, Yang et al., 2004; Kahn and Aubry, 2006) may be compared to osteo-

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porosis in vertebrates, as proposed by Stanley and Hardie (1998). Malformed coccoliths have lost their otherwise characteristic symmetry and smooth contours, and show all signs of anomalous (hypo- or hyper) calcification, some of which may be related to imbalance in calcium and magnesium concentrations (Herfort et al., 2004). What Stanley and Hardie (1998) describe as a failure in the late Neogene nannoplankton is in fact a change in morphological strategy. The calcareous nannoplankton originated during the Late Triassic, and by the Early Jurassic had radiated so extensively as to produce over half of the families that would diversify during the Mesozoic (Bown et al., 2004). They evolved and underwent their most significant radiation in Aragonite II sea (Early Jurassic). The coccolithophores thus became a hypercalcifying group not because they evolved as such, but because they had the physiologic adaptability to thrive in both aragonitic and calcitic seas. In Calcite II sea they produced heavily calcified scales, among which the most prominent may be Nannoconus. However, the coccolithophores were able to physiologically adapt to the transition from Calcite II sea to Aragonite III sea, and maintain a high diversity, but (and essentially if we consider the net contribution of coccoliths to sedimentation) losing the preponderant role of a hypercalcifier that they had acquired during Calcite II. CONCLUSIONS By examining the extant coccolithophores from the vantage point of the paleontologic record, I have shown that, despite heterogeneous phylogenetic composition, they are remarkably uniform with regard to size and structure of coccoliths and coccospheres, uniformity that results from morphologic convergence toward a specific strategy. In addition to selecting for small size of coccoliths and coccospheres, and abundant production of coccoliths, their morphological strategy probably favors polymorphism, dithecatism and an enhanced life cycle. I have shown that this global strategy (almost all taxa, and at all depths in the photic zone) was established by the earliest Pleistocene, and resulted from a long-term, progressive Neogene trend that began early in the Miocene and an accelerated Pliocene turnover between 3.7 and 1.95 Ma. The long-term trend toward smaller coccospheres with increasingly numerous, smaller, and simpler (lighter) coccoliths involved not only the rise of similar innovations in different lineages and habitats, but also the loss of large and complex morphostructures. The mid-Pliocene turnover involved both the loss of genera (morphostructural groups) that had been successful through the Paleogene and Miocene and morphologic innovations in the Noelaerhabdaceae. I have shown that the long-term trend toward small coccolith size parallels, and is likely related to, the rise of the Mg/Ca ratio and concomitant decrease in Ca+2 concentration, although this may not explain the decrease in cell size. I have proposed that the dominance and radiation of the small Noelaerhabdaceae since the middle Pliocene is linked to climatic variability. Although the small size strategy was well established by the earliest Pleistocene, the compositions of early

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Pleistocene and extant communities of coccolithophores are only remotely similar with regard to species composition. In fact, the Recent physiognomy of the coccolithophores cannot be extended beyond 0.050 Ma, when E. huxleyi became dominant. This implies that morphologic strategies are not static. Evolutionary processes occur within the limits imposed by a global strategy. From a methodological point of view, I have shown that coccospheres provide unique and necessary evidence that is complementary to the information gathered from the coccoliths. Even if scarce in the fossil record, they help us understand the paleobiologic dynamics of a group whose skeletons are structurally so exquisite as to reveal much of its history, thereby providing direct insight into the evolution of the earth system. I have also explained the limits of using extant species for reconstructing past changes in the earth system. In conclusion, this study has shown that while much attention is given to the massive biotic events that have been associated with dramatic terrestrial or extraterrestrial phenomena, subtle perturbations of the earth system have been no less determinant in fashioning the evolution of biological communities. Planktonic foraminifera and molluscans have also suffered significant Pliocene extinctions (Berggren et al., 1995b; Stanley and Campbell, 1981; Jackson et al., 1993). It will be interesting to determine whether these were associated with changes in morphologic strategy as in the calcareous nannoplankton. ACKNOWLEDGMENTS This paper was presented at the symposium “Mass Extinctions and Other Large Ecosystem Perturbations: Extraterrestrial and Terrestrial Causes” organized by Michael Rampino, Simonetta Monechi, and Rodolfo Coccioni at the 32nd International Geological Congress in Florence. I am grateful to them for editing this GSA Special Paper and to the editorial staff at GSA. My warm thanks go to Luc Beaufort, William A. Berggren, Paul Falkowski, Zoe Finkel, Alicia Kahn, Tara Kheradyar, Maria Triantaphyllou, and Colomban de Vargas for discussion on various aspects of this research, and to Anne Racciopi and Jack Cook for drafting the figures. I am especially grateful to Luc Beaufort and an anonymous reviewer for their thoughtful reviews of the manuscript. This study was partly supported by National Science Foundation (NSF) grants OCE 00 84032 Biocomplexity: The Evolution and the Radiation of Eukaryotic Phytoplankton Taxa (EREUPT) and DEB 0415351: Species-level Diversity and Evolution in Phytoplanktonic Coccolithophores Using Molecular, Morphological, and Fossil Data, and by a grant in support of micropaleontologic studies by PDVSA (Petróleos de Venezuela SA), Venezuela. REFERENCES CITED Aubry, M.-P., 1984, Handbook of Paleogene calcareous nannofossils, Volume 1: New York, Micropress, American Museum of Natural History, 266 p. Aubry, M.-P., 1988a, Phylogeny of the Cenozoic calcareous nannoplankton genus Helicosphaera: Paleobiology, v. 14, p. 64–80. Aubry, M.-P., 1988b, Handbook of Paleogene calcareous nannofossils, Volume 2: New York, Micropress, American Museum of Natural History, 279 p.

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Printed in the USA

The Geological Society of America Special Paper 424 2007

The Paleocene-Eocene Thermal Maximum in Egypt and Jordan: An overview of the planktic foraminiferal record Elisa Guasti* Department of Geosciences (FB 5), Bremen University, Post Office Box 330440, 28334 Bremen, Germany Robert P. Speijer Department of Geography and Geology, Katholieke Universiteit Leuven, Celestijnenlaan, 200E, 3001 Leuven, Belgium

ABSTRACT In the present study, we investigate upper Paleocene to lower Eocene planktic foraminiferal assemblages in Egypt and Jordan across a middle neritic to upper bathyal transect of the Tethyan continental margin. In particular, we evaluate the planktic foraminiferal turnover across the Paleocene-Eocene Thermal Maximum (PETM). Dissolution affects the planktic assemblages more intensively than previously considered, especially in the marls below the PETM. High numbers of Subbotina, fluctuating planktic/benthic (P/B) ratios, and low numbers of planktic foraminifera per gram (PFN) are indicative of dissolution, probably as a consequence of deep weathering. Hence, high numbers of Subbotina in this area do not indicate cooling. Despite this taphonomic overprint, we observe that well-diversified planktic foraminiferal assemblages of Subzone P5a abruptly changed into oligotaxic assemblages dominated by Acarinina during the PETM. Because various biotic and geochemical proxies indicate increased nutrient supply to the basin, we argue that the blooming of Acarinina is not indicative of oligotrophic conditions. Instead, we postulate that (low-trochospiral) Acarinina may have been better adapted to thrive under stressed environmental conditions, possibly because they hosted symbionts different from those in Morozovella. Keywords: PETM, planktic foraminifera, excursion taxa, Middle East. INTRODUCTION

(Aubry, 1998), diatoms (Oreshkina and Oberhänsli, 2003), and larger foraminifera (Scheibner et al., 2005). Accompanying this interval is a negative 2‰–3‰ carbonate isotopic excursion (CIE) (Kennett and Stott, 1991; Koch et al., 1995). What triggered the event is still under debate. The most widely accepted idea is that initial deep-sea warming led to the massive dissociation of oceanic methane hydrates, leading to further warming (Dickens et al., 1995). Other theories involve an increase of volcanic emission (Eldholm and Thomas, 1993), cometary impact (Kent et al., 2003), or intrusion of mantle-derived melts into carbon-rich sediments in the northeast Atlantic (Svensen et al., 2004).

The Paleocene-Eocene Thermal Maximum (PETM) represents a period of extreme global warmth (Zachos et al., 1993), associated with a major extinction of deep-sea benthic foraminifera (Tjalsma and Lohmann, 1983; Thomas, 1998) and evolutionary rejuvenations among planktic foraminifera (Kelly et al., 1996b), mammals (Clyde and Gingerich, 1998), calcareous nannofossils *Present address: TNO-Geobiology Team, Princetolaan 6, NL-3584 CB Utrecht, The Netherlands; [email protected].

Guasti, E., and Speijer, R.P., 2007, The Paleocene-Eocene Thermal Maximum in Egypt and Jordan: An overview of the planktic foraminiferal record, in Monechi, S., Coccioni, R., and Rampino, M.R., eds., Large Ecosystem Perturbations: Causes and Consequences: Geological Society of America Special Paper 424, p. 53–67, doi: 10.1130/2007.2424(03). For permission to copy, contact [email protected]. ©2007 The Geological Society of America. All rights reserved.

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It is also heavily debated whether this “supergreenhouse” period was generally associated with increased or decreased oceanic productivity. Several oceanic records and particularly the calcareous nannofossils suggest widespread oligotrophy (Kelly et al., 1996b; Bralower, 2002). In contrast, most continental-margin records suggest an increase in productivity (Speijer et al., 1996b, 1997; Schmitz et al., 1997b; Crouch et al., 2001; Speijer and Wagner, 2002; Gavrilov et al., 2003) and also some open-ocean records provide evidence for this (e.g., Thompson and Schmitz, 1997; Bains et al., 2000; Thomas et al., 2000; Stoll and Bains, 2003). An overall increase in oceanic productivity and burial of organic carbon in marine sediments would provide an important negative feedback mechanism (Bains et al., 2000; Dickens, 2001; Speijer and Wagner, 2002). Bains et al. (2000) suggested that an increase in nutrient supply may have caused blooms in marine phytoplankton, thereby sequestering the greenhouse gas CO2 to the deep sea by ~60 k.y. of enhanced biological pumping. This bloom might have been a response to a combination of increased weathering and runoff from the continents, oceanic fertilization from volcanic fallout, rising temperatures, and increasing atmospheric CO2 concentrations (Bains et al., 2000, and references therein). The southern Tethyan margin is of particular interest for studying lower Paleogene continental-margin records, as it provides continuously well-exposed outcrops and well-preserved material for micropaleontological research. For these reasons, the global boundary stratotype section and point (GSSP) for the Paleocene-Eocene boundary has recently been defined in Egypt within the Dababiya Quarry section, near Luxor in the Nile Valley (Ouda and Aubry, 2003). The PETM in Egypt has been intensively investigated mainly on the basis of smaller benthic foraminifera (Speijer, 1994; Speijer et al., 1996a, 1996b, 1997; Youssef, 2004; Alegret et al., 2005; Ernst et al., 2006), larger benthic foraminifera (Scheibner et al., 2005), planktic foraminifera (Obaidalla, 2000; Berggren and Ouda, 2003; Ouda et al., 2003), nannoplankton (Aubry et al., 1996; Monechi et al., 2000, Tantawy et al., 2003; Youssef, 2004), ostracodes (Speijer and Morsi, 2002; Elewa and Morsi, 2004), and geochemical and mineralogical parameters (Charisi and Schmitz, 1995, 1998; Schmitz et al., 1996, 1997b; Bolle et al., 2000; Speijer and Wagner, 2002; Dupuis et al., 2003; Knox et al., 2003). In this study, we evaluate planktic foraminifera assemblages, along a middle neritic to upper bathyal transect, from exposures in Egypt (Nile Valley, Eastern Desert, and Sinai) and in Jordan (Fig. 1). They are situated in an extensive epicontinental basin that covered most of Egypt, Israel, and Jordan at the end of the Paleocene (Speijer and Wagner, 2002), generally deepening in a northwest direction. The area is characterized by two major tectonic provinces, the stable shelf in the south (also known as the Nile Basin) and the unstable shelf in the north (Syrian Arc) (Said, 1990; Shahar, 1994; Tawarados, 2001). The localities studied in this work belong to the stable shelf (Said, 1990). Benthic foraminiferal data indicate that shallowest deposition occurred at Gebel Duwi (middle neritic, ~50–100 m) and the deepest at Wadi Nukhl (upper bathyal, ~500–600 m) (Speijer et al., 2000). Gebels

N

Mediterranean Sea Amman

Jordan Gebel Qurtayssiat

Cairo

30°N

Wadi Nukhl

Egypt

500 m

300 m Gebel Qreiya

26°N

Dababiya 0

Gebel Aweina

200 km 100 m 32°E

Gebel Duwi

Red Sea 36°E

Source: GEBCO.

Figure 1. Location map of the studied profiles (black dots). During the Neogene, the Jordanian localities shifted ~100 km north in response to sinistral movements along the Dead Sea transform fault (Garfunkel and Ben-Avraham, 1996). The normal dashed lines indicate the estimated paleobathymetry based on Speijer and Van der Zwaan (1994). The thick dashed line separates the stable shelf (south) from the unstable (north part).

Aweina, Qreiya, Dababiya, and Qurtayssiat represent outer neritic localities with paleodepths between 150 and 200 m. This transect provides a good opportunity to document the planktic foraminiferal developments across the PETM. Planktic foraminifera assemblages dominated by Acarinina characterize the PETM in open-marine environments worldwide. This has been interpreted as an indication of oceanic oligotrophy (Kelly et al., 1996b, 1998). However, we consider it highly unlikely that oligotrophic conditions of the open ocean could also prevail on the Tethyan margin, where, by contrast, all other proxies suggest an increase of nutrients and/or productivity. Therefore, we reevaluate previous paleoecological interpretations of this typical PETM planktic foraminifera assemblage. MATERIAL AND METHODS We collected samples at six locations. For Gebel Qreiya, Wadi Nukhl, and Gebel Qurtayssiat we used low-resolution sample sets (~1 m). For Duwi and Aweina centimeter scale sample sets across the PETM were available. Samples of the Dababiya DBH section, a 6-m-thick section comprising the GSSP of the Paleocene-Eocene boundary, were kindly provided by Christian Dupuis (Polytechnic of Mons). In order to further enhance the resolution of the Dababiya DBH record, we collected additional samples on the occasion of the inauguration of the GSSP of the P-E boundary during the Climate and Biota of the Early Paleo-

Paleocene-Eocene Thermal Maximum in Egypt and Jordan gene conference (CBEP IV) in Luxor, 2004. In the Dababiya and Qreiya sections, the P-E boundary is well defined at the basis of a non-calcareous clay to siltstone (Dupuis et al., 2003; Knox et al., 2003). In the Aweina section, the P-E boundary is situated at an omission surface between the shaley marls of Esna unit 1 and the overlying more limey (“calcarenitic”) bed. Abundant bioturbations extend some 7 cm down from the latter bed into the underlying shaley marls, indicating the absence of the lower PETM beds (Speijer et al., 1996a; Schmitz et al., 1997b). All samples were processed according to standard micropaleontological procedures explained by Speijer et al. (1996b) (samples from Egypt) and by Guasti (2005) (samples from Jordan). The fraction >125µm was used for all foraminiferal studies. We determined compositional data of the planktic foraminiferal assemblages, counting 200–300 specimens classified at genus level, generally using the concepts of Berggren and Norris (1997) and Olsson et al. (1999). In addition, we calculated the percentage of planktics in the foraminiferal association (planktic/benthic, or P/B, ratio, expressed as 100xP/(P+B) (cf. Van der Zwaan et al., 1990), as well as the total number of planktic foraminifera per gram of dry sediment (PFN) were calculated in order to identify taphonomic alteration of the foraminiferal assemblages. Carbonate content and whole-rock stable isotopes (O, C) were measured for the Qurtayssiat section at the Free University of Amsterdam (see Table 1). These parameters have been presented for the other sections in earlier studies (Speijer et al., 1996a; Schmitz et al., 1996; Speijer and Wagner, 2002; Dupuis et al., 2003). STRATIGRAPHY Lithostratigraphy In the Nile Valley and Eastern Desert (Egypt) the upper Paleocene to lower Eocene marls and shales belong to the Esna Formation, which is intercalated between the limestones of the upper Paleocene Tarawan Formation and the lower Eocene Thebes Formation (Said, 1990). In Sinai, the Tarawan Formation is often absent; therefore, the entire Paleocene to lower Eocene succession is often referred to as the Esna Formation (see details in Scheibner et al., 2001). Stratigraphy of the Duwi, Aweina, and Qreiya sections were discussed extensively by Speijer et al. (2000), the Dababiya section by Dupuis et al. (2003). We tenta-

TABLE 1. CARBONATE CONTENT AND CARBONATE ISOTOPIC RECORD FROM GEBEL QURTAYSSIAT 13 δ C Profile Sample Level CaCO3 (%) (m) (‰) Gebel Qurtayssiat, JQ 68 64 75.17 –2.14 67 63.2 48.86 0.29 66 62.5 56.94 –0.27 65 61.9 50.01 –1.92 64 61.1 39.54 –2.19 63 60.5 56.34 –0.18 62 59.8 60.81 1.11

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tively adopt the lithostratigraphic subdivision of the lower part of the Esna Formation as proposed by Dupuis et al. (2003) for the Dababiya section and applied to the Qreiya section by Knox et al. (2003) (Fig. 2). This subdivision consists of Esna unit 1, ranging from the top of the Tarawan Formation to the base of the PETM beds, known in the Nile Valley as the Dababiya Quarry Beds. Esna unit 1 consists of gray shaley marls. Esna unit 2 comprises the Dababiya Quarry beds and the overlying gray shales and marly shales. In full expression, the Dababiya Quarry Beds are composed of a succession of five different beds with an upward-increasing carbonate content. At the base, DQB1 is a dark, laminated non-calcareous clay without calcareous foraminifera, whereas at the top, DQB5 is a calcarenitic marl, largely composed of foraminifera. The middle part of the Dababiya Quarry beds, particularly DQB3 and the top of DQB2, is rich in apatite, largely in the form of coprolites and fish remains. The Paleocene-Eocene boundary coincides with the base of the clay bed of DQB1 (1.56 m above the base of the DBH section) at Dababiya (Dupuis et al., 2003). In addition to the sharp contact from marl to shale, it is characterized by a pronounced shift in δ13C of organic carbon, thus providing an excellent tool for correlation with deep-sea and terrestrial records spanning the P-E transition. The bio- and chemostratigraphy of these Egyptian localities were documented by Speijer et al. (2000) and Speijer and Wagner (2002). These authors pointed out that at Qreiya and Wadi Nukhl a dark-brown marl bed rich in total organic carbon (TOC), coprolites, fish bones and scales, and foraminifera marks the onset of the PETM. According to the criteria of Wignall (1994) this bed qualifies as black shale and was denoted as such. This bed was not observed at Aweina, thus pointing to an unconformity across the P-E boundary in this section. Evaluation of results from Egypt in the context of the detailed Dababiya and Qreiya sequence (Dupuis et al., 2003; Knox et al., 2003) indicates that the coprolite-rich, dark-brown marl bed (“black shale”) observed at Qreiya (Speijer et al., 2000; Speijer and Wagner 2002) corresponds to DQB3. Consequently, this bed does not correspond to the true onset of the PETM, as considered earlier, but rather to the level with minimum δ13C values and thus a somewhat later stage during the PETM (Dupuis et al., 2003). In Wadi Nukhl, Qreiya, and Qurtayssiat the low-resolution sampling might be the main cause for not identifying all DBQ beds in our records (mainly DBQ1–2 are absent from our data set). Recent studies on Qreiya by (Knox et al., 2003; Berggren and Ouda 2003; Soliman 2003) and new field observations indicate that this is true at least for the Qreiya section. Over a large area in Egypt, DQB5 is the most distinct bed within the relatively monotonous Esna succession, especially in weathered outcrops, because in all sections it has anomalously high CaCO3 content (50%–70%) in the lower part of the Esna Formation. This bed corresponds to the bed previously denoted as Esna unit C at Aweina (Speijer et al., 1996b) and more generally as the “calcarenitic bed” (Schmitz et al., 1997b; Speijer et al., 2000) because of its enormous abundance of planktic foraminifera (>10,000 specimens/gram in the size fraction >125

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Figure 2. Stratigraphic correlation of the localities, arranged along a transect from the deeper localities on the left to the shallowest ones on the right (roughly a N-S orientation). Planktic foraminiferal biostratigraphy follows Speijer et al. (2000); Gebel Aweina has been vertically exaggerated. DQB—Dababiya Quarry beds.

microns). This bed records the return to more positive δ13C values (Schmitz et al., 1997b; Speijer et al., 2000; Dupuis et al., 2003). At Aweina, this equivalent to DQB5 directly overlies Esna 1. Thus DQB1–4 were either not deposited or not preserved at Aweina, marking a significant hiatus at the P-E boundary (Speijer et al., 2000). Ouda et al. (2003) provided a strongly deviating interpretation of the stratigraphic sequence across the P-E boundary at Aweina. These authors suggested that a 1-m-thick interval below the calcarenitic bed records the onset of the carbon isotope excursion (CIE). These beds would thus correlate with the lower Dababiya Quarry Beds and instead the upper ones would be missing. The isotopic data consulted for this view were derived from the data presented by Charisi and Schmitz (1995) and Schmitz et al. (1996). However, these and other authors, including Ouda (2003), convincingly demonstrated that δ13C records based both on whole rock and Lenticulina spp. show virtually straight verti-

cal lines up to the unconformity, i.e., the base of the calcarenitic bed (Schmitz et al., 1996, 1997b; Speijer et al., 1996a, 2000; see also Figs. 2, 4, and 7, in Ouda, 2003). These values, ~1‰ for whole rock and around −1‰ for Lenticulina spp., fit well within the range of pre-PETM values elsewhere in the region and are much higher than values (−1‰ to −2‰) for whole-rock measurements found within the PETM in the region (e.g., Speijer et al., 2000). If the view of Ouda et al. (2003) were correct, the Aweina record should show a negative excursion below the calcarenitic bed (DQB5), followed by an abrupt positive shift at the unconformity. The opposite in fact is true: there is no shift below the unconformity and a small negative shift above it. Biostratigraphic data, such as the presence of Acarinina sibaiyaensis and A. africana just below the unconformity (Ouda et al., 2003), cannot be considered as evidence for the opposite view, because these taxa were originally described from the top of the Tarawan Formation and from Esna unit 1 (El-Naggar, 1966; Speijer et al., 2000). Thus

Paleocene-Eocene Thermal Maximum in Egypt and Jordan we are confident that the deposits immediately underlying DQB5 at Aweina constitute the top of the Paleocene of Esna 1 unit, not the base of the Eocene. Consequently also the correlation of the P-E boundary at Aweina to other Nile Valley sections, such as G. Kilabiya, G. Qreiya, and Abu Ghurra (Berggren et al., 2003; Ouda and Berggren, 2003; Ouda et al., 2003), needs revision. In Gebel Duwi, the lithostratigraphic equivalent to DQB3 developed somewhat differently from Dababiya and Qreiya as a ~20-cm-thick fissile pink marl, containing abundant phosphatic peloids and fish remains (Speijer et al., 2000). Underneath this bed, Speijer et al. (2000) observed a 1-cm-thick shale bed without foraminifera, overlying a 20-cm-thick reddish interval with abundant calcitic pseudomorphs of dolomite rhombs and few foraminifera. It is not unlikely that this interval could in part correspond to DQB1–2. Similar to Gebel Aweina and Wadi Nukhl, a calcarenitic bed, correlative with DQB5, is present in the upper part of the CIE. From our data it is not clear whether an equivalent to DQB4 is present at Duwi or not. In Jordan the studied interval belongs to the pale yellow, light-gray chalky and marly sediments, which have been named Muwaqqar Chalk Marl Formation (MCM) by Masri (1963). Like the Esna Formation in Egypt, the MCM is rich in secondary gypsum veins. It is overlain by the chert-rich chalky limestones of the Umm Rijam Chert Limestone Formation. Within the monotonous marly MCM sediments, a ~1-m-thick dark-purple bed, partially laminated and rich in fish remains, is intercalated. Lithological and sedimentological similarities and biostratigraphic data suggest that this bed correlates with DQB3.

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Speijer et al., 2000; Norris and Nunes, 2004; Guasti, 2005). This results mainly from taxonomic ambiguities, but these need to be resolved prior to introducing a formal subzonation of Zone P5. Hence, until these ambiguities are clarified we prefer the use of Subzone P5b based on the occurrence of M. allisonensis, which thus far has been observed exclusively within the PETM (Kelly et al., 1996b, 1998; Speijer et al., 2000). Subzone P5b marks DQB3 in every locality. Within the high-resolution context of Dababiya this subzone represents the middle part of the CIE, i.e., the level where δ13C values reach minimum values and start to increase again to a level midway the gradual return to stable values. In Gebel Qurtayssiat, the black shale (DQB3) contains a calcareous nannofossil assemblage indicative of CP8b of Okada and Bukry (1980), based on the first occurrence of Rhomboaster (in particular R. calcitrapa/bitrifida, R. cuspis, and R. spineus). This bed immediately overlies marls corresponding to Zone CP4, indicating the presence of an expanded hiatus (Eliana Fornaciari, 2004, personal commun.). The planktic foraminiferal assemblages, just below the PETM, record an interval of strong dissolution (probably in Zone P4). RESULTS Planktic Foraminiferal Assemblages

Biostratigraphy

The planktic foraminiferal assemblages in the study area are composed mainly of Acarinina, Morozovella, and Subbotina, which together made up >80% (Fig. 3). Igorina, Parasubbotina, and Globanomalina are generally <5%. Chiloguembelina and Zeauvigerina always occur in background numbers (<1%).

The studied interval belongs to planktic foraminifera Zone P5 (Berggren et al., 1995), defined as the biostratigraphic interval between the Last Appearance Datum (LAD) of Globanomalina pseudomenardii and the LAD of Morozovella velascoensis. We apply a subdivision (Subzones P5a–P5c) proposed by Speijer et al. (2000), which consists of three subzones: Subzone P5a or Globanomalina chapmani Interval Subzone, Subzone P5b or Morozovella allisonensis Total Range Subzone, and Subzone P5c or the Globanomalina luxorensis Interval Subzone. Subzone P5b was originally considered to represent the early part of the PETM. Morozovella allisonensis is common and well preserved in deposits equivalent to DQB3 at Qreiya and Gebel Duwi. We have also observed it in our samples from Nukhl, Dababiya, and Qurtayssiat, in DQB3, though it is not common. According to Berggren et al. (2000) and Berggren and Ouda (2003), the application of this scheme has been problematic, because M. allisonensis occasionally occurs too sporadically. Instead, Berggren and Ouda (2003) favored the range of the more common A. sibaiyaensis to characterize a subzone correlative with the PETM. However, it has been demonstrated that the subbiozonation of Zone P5 based on the occurrence of Acarinina sibaiyaensis leads to miscorrelations, because this taxon, together with A. africana, occurs well before the PETM (El-Naggar 1966;

Pre-PETM In the marls below the PETM, Subbotina exhibits maximum values: 13%–27% at Duwi, ~30% at Dababiya, and up to ~40% at Aweina and Wadi Nukhl. A maximum of 62% is recorded at Qreiya. At Nukhl, Globanomalina is present in higher numbers (~7%) compared to the other localities. Among the surface dwellers, Morozovella is the most abundant genus and it ranges from a minimum value at Qreiya (12%), ~40% at Nukhl, to a maximum at Duwi (~60%). At Duwi, the abundance of Igorina is slightly higher than in the others localities (up to 8%). Acarinina makes up ~27% in Duwi and Qreiya and ~20% in Nukhl. At Aweina a maximum abundance of Acarinina (~50%) occurs just underneath the unconformity at the P-E boundary, where Morozovella decreases to 28%. Striking is the common occurrence of A. sibaiyaensis and A. africana in the uppermost 50 cm of the Paleocene at Aweina. P/B ratios, expressed as %P, generally fluctuate between 35% and 75% at Aweina and Dababiya (Fig. 4). Just underneath the unconformity at the P-E boundary at Aweina, a peak of 97% is sandwiched between two lower values (45% below and 25% above). The values are more constant at Nukhl (~75%) and Duwi (~65%). At Qreiya a minimum value (6%) is recorded at the base of Esna unit 1.

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Figure 3. Relative frequency of planktic foraminiferal assemblages. Planktic foraminiferal biostratigraphy follows Speijer et al. (2000); Gebel Aweina has been vertically exaggerated.

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Figure 4. Planktic/benthic ratio (%P), the numbers of planktic foraminifera/gram (PFN) expressed as logarithmic scale and the carbonate content (CaCO3%). Planktic foraminiferal biostratigraphy follows Speijer et al. (2000); Gebel Aweina has been vertically exaggerated. CIE indicates the carbonate isotopic excursion.

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Planktic foraminiferal numbers (PFN) range between 150 and 4500 at Duwi (Fig. 4). At Aweina, Nukhl, and Dababiya these numbers are lower, between 20 and 300. At Qreiya the lowest value is recorded at the base of Esna unit 1; only five planktic specimens per gram. Pre-PETM beds contain 40%–50% CaCO3 at Duwi, Dababiya, and Nukhl, and ~30% in Aweina. Just below the PETM, the values decrease to ~35% at Nukhl and 20% at Qreiya. From these data we observe that depressed P/B ratios, lower PFNs, and low carbonate contents correspond with high numbers of Subbotina, in samples relatively rich in oxidized pyritic burrow fills. PETM A major change in the planktic foraminiferal assemblages occurs within the lower part of the PETM. In beds DQB1–2 at Dababiya, no planktic foraminifera are present, except for a relative peak of poorly preserved Acarinina (~90%) in sample DBH 2.3 (bed DQB2), in which a multichambered variety of Acarinina sibaiyaensis also is recorded. At Duwi these beds possibly correspond to 1 cm shale and the underlying dolomitic interval (~20 cm). In the other localities these beds are not recorded in our sample sets. Instead, DQB3 and equivalent beds can be traced in all sections except Aweina and are characterized by very high numbers of Acarinina (~80%). In Qurtayssiat, Acarinina is slightly less abundant (~60%), and Subbotina, Parasubbotina, and Globanomalina exhibit higher values (30%, 7%, and 6%, respectively), compared to the other sections. In this interval M. allisonensis and a multichambered variety of Acarinina sibaiyaensis are recorded, of which the latter is the major component. Acarinina sibaiyaensis (sensu El-Naggar, 1966) and A. africana also occur, but in low numbers. At Dababiya in sample DBH 3.12 Acarinina decreases (~30%) and Parasubbotina reaches a maximum (~21%). Maximum P/B ratios (>99%) are encountered at Wadi Nukhl, Duwi, and Qreiya, whereas in Gebel Qurtayssiat the P/B ratios decrease (~64%). In DQB2, just one sample, DBH 2.3, contains enough foraminifera to have a reliable P/B ratio: its P/B ratio is ~100%, whereas sample DBH 2.72 contains only ten planktic foraminifera in the whole residue. Planktic foraminiferal numbers of the lower PETM beds vary between the sections, but they are generally higher than in the marls below. At Dababiya these values are <10 in DQB2–lower DQB3 and they increase up to ~100 in upper DQB3. These numbers are up to 3700 at Duwi, ~1400 at G. Qreiya, and ~4300 at Wadi Nukhl. Only at Qurtayssiat the value is extremely low (~2/g). At Dababiya, at the base of DQB1 the carbonate content drops to zero, increasing from the lower part of DQB2 upward. From sample DBH 2.5 onward the values increase again to 30%. In DQB3 and equivalents at Dababiya, Duwi, and Qurtayssiat, the carbonate content is generally ~40%, and ~50% at Qreiya and Wadi Nukhl. In DQB4–5, above the interval of Acarinina dominance, the planktic foraminiferal assemblages become more diversified, as indicated by the gradual decrease of Acarinina, and the increase

of the other genera. Specimens of A. sibaiyaensis and A. africana are rarely found in DQB4–5 at Dababiya and Aweina, and neither M. allisonensis nor a multichambered variety of Acarinina sibaiyaensis occur anymore. Overall, Morozovella ranges from 36% to 46% and Subbotina increases, but the values remain lower than in Subzone P5a (6%–17% at Duwi and Aweina). In Aweina and Duwi, Globanomalina is more common (<4% and <6%, respectively) than in Subzone P5a. The P/B ratios in DQB4–5 increase up to 80%–95% and the values are fairly constant in each locality. Similarly, the PFNs abruptly increase, between 10,000 (at Aweina) and 15,000 (at Dababiya and Duwi). At Qreiya the numbers are lower, ~1200 (DQB4). Also the carbonate contents increase in DQB4–5, ranging from 50% at Aweina, to >60% at Dababiya and >80% at Nukhl. Post-PETM Above DQB5, in the post-PETM phase, planktic foraminiferal assemblages continue to diversify in all localities. At Dababiya and Nukhl, Morozovella decreases toward the uppermost part with minimum values on top (21% and 12%, respectively), whereas Subbotina increases (up to 42% in both localities). An opposite trend is observed in Qreiya, where Subbotina decreases and Morozovella increases upward. Numbers of Acarinina are ~30% in Nukhl and Dababiya, higher in Qreiya and Duwi (~40%), and 40% or more in Gebel Qurtayssiat. Generally the P/B ratios are quite stable between 85% and 90%, slightly higher in Nukhl (>90%). The numbers of foraminifera are generally quite stable ~1000/g, with a maximum peak in Duwi (~18,000/g) and a minimum in Nukhl (~90/g). Generally, the carbonate content values are quite stable around 30%, but in Nukhl they are higher (~50%). At Duwi, on the other hand, from sample 1033 onward, an interval with lower carbonate content (~20%) is recorded, interrupted by a ~1.5-m-thick interval with no carbonate between samples 1040 and 1042. DISCUSSION Avoiding Taphonomic Pitfalls The paleoecological significance of foraminiferal assemblages is based on the assumption that the fossil assemblage is a good reflection of the original live assemblage. However, the original ecological assemblage is transformed by a number of processes, of which differential test production (population dynamics) and preservation (taphonomy) are the most important ones (Martin, 1993). Processes of population dynamics are difficult to reconstruct in past environments and especially for extinct taxa. Taphonomic distortion of the assemblages is more easily identified (see below) and often is linked to the depositional environment (Loubere and Gary, 1990; Martin, 1993). Despite the transformation from the live assemblage to the dead assemblage, the fossil foraminiferal assemblage still con-

Paleocene-Eocene Thermal Maximum in Egypt and Jordan tains important and detailed information of the (time-averaged) environmental conditions during deposition (e.g., Jorissen and Wittling, 1999; Murray and Alve, 1999). Dissolution of calcareous tests below the lysocline is a welldocumented and understood phenomenon in deep-sea oceanography and paleoceanography (e.g., Berger, 1970). Dissolution, however, is not exclusively restricted to the deep-sea. Partial or complete carbonate dissolution is also sometimes encountered in shallow marine deposits exposed on land. This process may occur during deposition in undersaturated (e.g., brackish) waters, during diagenetic alteration (e.g., with the formation of carbonate concretions), and under the influence of weathering (El Kammar and El Kammar, 1996). In addition, the metabolic consumption of organic carbon from benthic organisms within sediments may contribute to dissolution of calcite even above the lysocline (Freiwald, 1995; Dittert et al., 1999). Dissolution may also occur during sample processing in the laboratory, for instance if sediment containing pyrite-filled tests is disintegrated with the aid of an H2O2 solution. Besides the potentially mechanically destructive bubbling from an H2O2 inside foraminifera tests (e.g., Hodgkinson, 1991), dissolution can be a serious problem: the rapid oxidation of pyrite leads to the formation of sulfuric acid, lowering the pH of the solution within the tests. In this way, pyritefilled tests will partially or fully dissolve. The effect is similar to long-term weathering of pyrite-filled tests in outcrops. In studies on microfossil assemblages from continental-margin deposits, partial dissolution is an often ignored problem. The problem becomes apparent when, for instance, the P/B ratio is used as an indication for paleodepth and sea-level change. It is well known that tests of planktic foraminifera are generally more susceptible to dissolution than those of hyaline benthic foraminifera (Douglas and Woodruff, 1981). Moreover, in the modern ocean the more solution-susceptible planktic foraminifera species are relatively small in size and have large pores and thin walls, whereas the less solution-susceptible species have large tests with small pores and thick walls, and, in general, spinose species are less resistant than the non-spinose ones (Bé, 1977). Moreover, Paleocene planktic foraminifera appear to have dissolved differentially, as pointed out by Boersma and Premoli Silva (1983). These authors constructed a solution-susceptibility ranking for planktic genera, in environments presumably below the foraminiferal lysocline. In order of decreasing susceptibility: the juveniles of most groups, the large morozovellids, acarininids, small morozovellids, smooth-walled genera, Parasubbotina and Subbotina (spinose taxa). Dissolution during diagenesis is probably controlled more by wall thickness and it acts first on smaller and thinner forms than on more robust taxa, and thus morozovellids are the most resistant ones (Boersma and Premoli Silva, 1983). There are further ways to assess the amount of dissolution in foraminiferal assemblages. For instance, planktic/benthic ratios can be evaluated against the total number of foraminifera per gram of sediment (PFN). The amount of CaCO3 in the sediment provides additional information. Where fluctuations between

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these records coincide, e.g., a drop in P/B ratio together with a drop in PFN and CaCO3, it is likely that carbonate dissolution has occurred at some step during the generation of the fossil assemblage (e.g., Speijer and Schmitz, 1998). In exceptional situations this association of changes may indicate a real paleoenvironmental change. If so, in shelf settings this should be readily discernable from the composition of a well-preserved planktic and/or benthic foraminiferal assemblage. Moreover, weathering in arid regions deeply affects the composition of the sediments. El Kammar and El Kammar (1996), in their study on the shales and marls of the Campanian to Paleocene Duwi and Dakhla Formations in Egypt, established that the total organic carbon (TOC) of these units diminishes dramatically upon weathering (from an average 7.05% to 0.04%). At the same time, the content of calcite is also drastically reduced by ~60%, owing to the carbonic acid produced as organic matter decays during weathering. Therefore, the increase of clay content is only apparent, deriving probably from a depletion of carbonate and organic matter contents. El Kammar and El Kammar (1996) also showed that fresh TOC-rich samples of the Dakhla Formation in the Eastern Desert could only be obtained from caves, not from normal exposures prone to weathering. Speijer and Wagner (2002) and Speijer (2003) measured TOC contents in the Esna and Dakhla Formations, respectively. Beyond the PETM, the TOC content of the Esna Formation is invariably low (0%–0.2%), whereas the black shales of the PETM record a peak of TOC (1.5%–2.7%). Also the TOC content in the Dakhla Formation ranges between 0.1% and 0.3%, and is similar to the outcrop values of El Kammar and El Kammar (1996), whereas the Danian/ Selandian black shale bed yields between 0.75% and 2.0% TOC. Considering that there is almost a complete loss of TOC due to weathering in Eastern Desert outcrops (El Kammar and El Kammar, 1996), we argue that the TOC contents measured by Speijer and Wagner (2002) and Speijer (2003) in outcrop samples are probably grossly underestimated. Speijer and Wagner (2002) could not exclude weathering as a cause of the low TOC values and the absence of organic dinoflagellate cysts, but they instead favored burial effects. The data of El Kammar and El Kammar (1996) demonstrate that weathering is more likely after all. Considering the amount of combusted organic carbon in the studied sediments and the conversion of all pyrite into iron-oxides, it is likely that this had an influence on the amount of carbonate and thus on the foraminiferal composition. The amount of gypsum veins in the studied marls provides another indication of the intensity of weathering as the gypsum is the by-product of pyrite weathering and dissolution of CaCO3. Indeed, the faunal composition also is indicative of different levels of dissolution. In the marls below the PETM of most of the studied sections (Dababiya, Aweina, Qreiya, and Nukhl), high numbers of Subbotina are often associated with large numbers of oxidized pyritic burrows and concretions. In addition, the P/B ratio, PFNs, and the carbonate content are generally lower than in the marls above the PETM. For instance, it is striking that the highest number of Subbotina is recorded at Qreiya, associated

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with the extreme low PFN and P/B ratio. We suggest that all these features together are an index of dissolution, which in general seems to have been more severe before the PETM than after. Therefore, deep weathering might be the principal factor responsible for dissolution in the marls below the PETM. However, we do not rule out that metabolic consumption of the organic carbon from benthic organisms also contributed to this pattern. Conversely, we could also conclude that originally there may have been less organic carbon in P5c sediments. However, not all localities seem to be affected by this phenomenon. For instance, at Duwi the percentage of Subbotina is generally lower than at the other localities and the carbonate content higher. Although these differences may primarily relate to the shallower setting, it could also be argued that the deposits at Duwi were less prone to weathering than in the Nile Valley. At Qurtayssiat, weathering also seems to have strongly affected the samples of Subzone P5b, as indicated by lower numbers of Acarinina, increased numbers of Subbotina, and altogether extremely low numbers of planktic foraminifera. In the uppermost Subzone P5c, Subbotina increases again and Morozovella decreases. Because P/B ratios, PFNs, and carbonate content are generally high, we assume that, in this case, the increased numbers of Subbotina represent a primary signal. This change in the planktic assemblage coincides with the termination of the PETM, so it may point to a termination of the extreme warming phase (Berggren and Ouda, 2003). However, superimposed on this general trend, high numbers of Subbotina are still related to lower P/B and PFN at certain levels, suggesting that probably dissolution occurred. In view of these considerations, we stress that assessments of foraminiferal assemblages must take into account that high numbers of Subbotina might be an artifact from differential dissolution, and are not necessarily related to ecological factors such as cooling. A combination of P/B ratios, PFNs, and carbonate content represents a good tool to discern whether increased Subbotina reflects a primary signal or not. Response of Planktic Foraminiferal Assemblages to the PETM In Zone P5, at each locality, Subbotina, Acarinina, and Morozovella are the major components of the planktic assemblages, whereas Igorina, Parasubbotina, and Globanomalina constitute just minor parts. The depth gradient ranging from middle neritic to upper bathyal deposits is not clearly expressed in a corresponding increase in P/B ratios from shallow to deep as could be anticipated from the basis of the relationship between modern P/B ratios and water depth (Van der Zwaan et al., 1990). One reason for this may be the configuration of relatively broad Paleogene shelves compared to those of today; another is that partial dissolution strongly overprints the P/B ratios. Despite this dissolution, it is possible to reconstruct general patterns of water-column conditions. Before the PETM, well-diversified planktic assemblages dominated by Subbotina, Acarinina, and

Morozovella are indicative of a stratified and well-oxygenated water column. The scarcity of Chiloguembelina, considered to be a low-oxygen thermocline dweller (Premoli Silva and Boersma, 1988; Boersma and Premoli Silva, 1989), supports this view. At Nukhl, higher numbers of Globanomalina are in agreement with the deeper setting of this locality and with colder water compared to the shallower localities. In contrast, at the shallowest Duwi site, the higher abundance of surface dwellers indicates a preference of these taxa for shallower and warmer conditions. In particular, increased numbers of Morozovella in the uppermost Subzone P5a are indicative of higher sea-surface temperatures (i.e., Norris, 1996; Berggren and Norris, 1997; Quillévéré and Norris, 2003). A major change in the assemblages occurs in the early to middle PETM. The DQB2–3 beds and their equivalents contain oligotaxic assemblages, dominated by generally poorly preserved Acarinina, suggesting high levels of environmental stress in the water column. In particular, the multichambered variety of Acarinina sibaiyaensis is the most common species. A. sibaiyaensis and A. africana are also present, but very rare. Morozovella is almost completely absent during this interval, except for M. allisonensis, which, like A. multicamerata, seemed to have evolved suddenly. In addition, the impoverishment of the thermocline dwellers (Subbotina and Parasubbotina) and the oligotaxic calcareous benthic foraminifera assemblages (Speijer et al., 1996b; Speijer and Wagner, 2002) suggest intensification and vertical expansion of the regional Oxygen Minimum Zone (OMZ). Above DQB3 and correlative beds (Subzone P5c) the number of Acarinina gradually decreases and the population becomes more diversified. In particular, various Acarinina species are present (i.e., A. sibaiyaensis, A. africana, A. coalingensis, A. soldadoensis). All other genera slowly increase again; in particular, Morozovella completely recovers in numbers and in species diversity. Whereas M. allisonensis disappears, the M. velascoensis and M. subbotinae groups return to become the most dominant taxa. Such diversified assemblages combined with an increased number of planktic foraminifera indicate recovery of water-column conditions after the PETM perturbation. Eutrophy versus oligotrophy during the PETM is still an open debate. It seems that oligotrophy characterized large parts of the deep ocean (Kelly et al., 1996b; Bralower, 2002), whereas along continental margins high-nutrient environments developed (Speijer et al., 1996b, 1997; Schmitz et al., 1997b; Crouch et al., 2001; Speijer and Wagner, 2002; Gavrilov et al., 2003). Surprisingly though, the response of planktic foraminiferal assemblages to the climatic perturbation is similar in both settings. Acarinina dominates the planktic assemblages in the open ocean (Kelly et al., 1996b, 1998; Kelly, 2002) and in deep and shallow Tethyan basins (Arenillas and Molina, 1997; Schmitz et al., 1997a; Lu et al., 1998; Arenillas et al., 1999; Molina et al., 1999; Pardo et al., 1999; Obaidalla, 2000; Berggren and Ouda, 2003; Ouda, 2003). Various authors (Kelly et al., 1996b, 1998; Arenillas and Molina, 1997; Kelly, 2002) explain this increase as a response to oligotrophic conditions during the PETM, where photosymbiosis

Paleocene-Eocene Thermal Maximum in Egypt and Jordan may facilitate Acarinina to thrive in low-nutrient waters. However, other proxies, such as benthic foraminifera (Speijer et al., 1996b; Speijer and Schmitz, 1998; Thomas et al., 2000), calcareous nannofossils (Monechi et al., 2000), siliceous plankton and organic dinocysts (Benjamini and Sheva, 1992; Crouch et al., 2001; Egger et al., 2003), barium, phosphate, clay minerals, and TOC records (Schmitz et al., 1997b; Bains et al., 2000; Bolle et al., 2000; Schmitz, 2000; Speijer and Wagner, 2002), rather indicate eutrophic conditions, related to increased upwelling and/or weathering and runoff, particularly in the Tethyan area. For instance, Bolle et al. (2000) suggested that in the southern Tethys, humid and warm conditions in the hinterland enhanced runoff, supplying plenty of nutrients to the water column, to prevent extreme oligotrophy in the basin. In addition, Speijer and Wagner (2002) proposed a paleoceanographic model for this area, suggesting that an inflow of less-oxygenated intermediate water into the epicontinental circulation, combined with intensified upwelling, resulted in severe anoxia on the seafloor. This process resulted in an expanded OMZ and led to the suppression of the thermocline dwellers, such as subbotinids. Although low productivity on the southern Tethyan margin is an unrealistic scenario, we believe that oligotrophy alone cannot be the primary factor controlling the dominance of Acarinina. Clearly, the trophic strategy or strategies of this group need to be investigated more closely. As pointed out by Houston and Huber (1998), the role of photosymbionts in modifying stable isotopic values is not fully understood in living planktic foraminifera; therefore, the application of stable isotopes to discriminate symbiotic and asymbiotic taxa in the fossil records is full of uncertainties and suppositions. Additionally, we believe that the relationship between planktic foraminifera and their symbionts during the Paleocene is more intricate than generally asserted; however, it probably represents the key to demystifying the success of Acarinina and the suppression of Morozovella at the PETM. It is well accepted that Morozovella and Acarinina share similar ecological preferences: they supposedly inhabited the mixed layer and carried algal symbionts (Pearson et al., 1993; D’Hondt et al., 1994; Kelly et al., 1996a; Norris, 1996). However, among Acarinina different species preferred different habitats within the mixed layer. For instance, the low-trochospiral early species of Acarinina (prior to 57.0 Ma) probably lived in the deeper, cooler, and more mesotrophic part of the mixed layer (Quillévéré and Norris, 2003). Additionally, Corfield and Norris (1998) suggested that significant carbon isotope differences may exist between different clades of Morozovella; for instance M. subbotinae occupied a slightly deeper, cooler, and more mesotrophic habitat than M. velascoensis (Quillévéré and Norris, 2003). Therefore, the M. subbotinae lineage and early species of Acarinina seem to have shared a similar habitat (Olsson et al., 1999; Quillévéré and Norris, 2003). Moreover, oxygen isotopic values indicate that M. allisonensis and a multichambered variety of A. sibaiyaensis encroached into deeper water during the PETM Kelly et al. (1996b, 1998). We consider it highly relevant that during the PETM, the population of Acarinina is composed mainly of low-trochospiral specimens with higher oxygen

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isotopic values, more similar to early Acarinina species. Similarly, at the Danian-Selandian transition in the same area, we observe a bloom of low-trochospiral Acarinina during the “Neoduwi event” (Guasti, 2005, Chap. 4), in relation to increased runoff. Also in the Ypresian of the northern Tethyan margin, a peak of Acarinina is associated with sapropelic deposits and related to increased productivity (Oberhänsli and Beniamovskii, 2000). From these examples, we conclude that recorded peaks of Acarinina are particularly connected to increased primary productivity in the Tethyan area during the lower Paleogene. Also, Kelly et al. (2005) described a peak of A. subsphaerica at Site 690 (Weddell Sea, near Antarctica) in connection to increased continental weathering/runoff, a relationship that disagrees with the supposed affinity of this taxon for oligotrophy. The importance of the partnership between planktic foraminifera and photosymbionts is amply documented for extant (i.e., Hemleben et al., 1989) and fossil taxa. An impairment of this relationship has been suggested as a cause for the extinction of certain taxa. For instance, Kelly et al. (2001) suggested that a progressive deterioration of symbiosis led to the gradual extinction of Morozovella velascoensis. Additionally, during the late middle Eocene, the extinction of the whole Morozovella lineage seems to be related to increased surface-water productivity and the deterioration of photosymbiotic partnership with algae (Wade, 2004). In the modern ocean, several species of planktic foraminifera are characterized by symbiotic associations. For instance, the endosymbiont of Orbulina universa, Globigerinoides ruber, and G. sacculifer is the dinoflagellate Gymnodinium béii (Gast and Caron, 1996; Rink et al., 1998; Wolf-Gladrow et al., 1999), whereas Globigerinella aequilateralis, Globigerina cristata, and G. falconensis host symbiotic chrysophycophytes (Rink et al., 1998). The amount of light availability plays an important role in the abundance and distribution of symbiont-bearing planktic foraminifera. In the northern California Current, for instance, symbiotic species are more abundant in less turbid (and nutrientpoor) offshore waters, whereas asymbiotic species dominate in high-turbidity (and nutrient-rich) waters close to the coast (Ortiz et al., 1995). However, symbiotic taxa can also reach high abundance in nutrient-rich waters where turbidity is low and availability of light is high (Ortiz et al., 1995). Furthermore, in the Caribbean Basin increased abundance of Globigerinoides ruber is associated with enhanced primary productivity (Schmuker and Schiebel, 2002). Hence, high numbers of symbiotic planktic foraminifera do not point exclusively to oligotrophic conditions; instead light availability is often the determining factor. We are aware that it is arduous to demonstrate analogies of host-symbiont relationships between extant and Paleocene planktic foraminifera. However, we aim to point out here that oversimplified concepts of trophic strategies of Paleocene taxa are currently inadequate to explain the distributional patterns associated with the PETM and other transient events. We assume that trophic strategies and relationships with symbionts were also complex among Paleocene planktic foraminifera. This is difficult to investigate but should not be excluded a priori. We could speculate, for example,

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that Morozovella and Acarinina might have hosted different symbionts. As with bleaching in marine environments, we speculate that the relationship between Morozovella and its symbionts may have been suppressed at low-middle latitudes during the PETM, in response to environmental stress conditions, including high seasurface temperature. By hosting different symbionts (or having the capability of changing symbionts), low-trochospiral Acarinina may have been better adapted to these conditions. Hence, a mechanism for preferentially eliminating symbionts might have been the major factor controlling the suppression of Morozovella during the PETM. As a consequence, Acarinina could also successfully dominate the oligotrophic open ocean. We cannot exclude a change in the structure of water masses from having played a role in this faunal change; however, such a scenario on a global scale seems unrealistic to us. Clearly, our ecological inferences are speculative and need further evaluation, but they underscore that the ecology of Paleocene planktic foraminifera is still poorly understood and that the relationships between foraminiferal distributions and biotic and physicochemical parameters, and between the surface dwellers with their symbionts, are probably much more complex than currently envisaged. CONCLUSIONS Irrespective of their paleobathymetric positions, the studied sections in the Middle East are characterized by similar planktic foraminiferal assemblages in planktic foraminifera Zone P5. Within this zone, changes in the assemblages recorded paleoclimatic and paleoceanographic variations, in particular connected to the hyperthermal event at the Paleocene-Eocene boundary: • In the marls below the PETM (Subzone P5a), the planktic foraminiferal assemblages are affected by dissolution, indicated by high numbers of Subbotina, fluctuating P/B ratios, and lower numbers of planktic foraminifera per gram of sediment. We stress that assessments of foraminiferal assemblages must take into account that high numbers of Subbotina may be an artifact from differential dissolution, and are not necessarily related to ecological factors such as cooling. A combination of P/B ratios, PFNs, and carbonate content represents a good tool to discern whether an increase of Subbotina reflects a primary signal or not. • The relatively well-diversified planktic assemblages in Subzone P5a are replaced by an oligotaxic Acarininadominated assemblage in response to environmental stress during the PETM. Because biotic and geochemical proxies indicate increased nutrient supply into the basin, due to upwelling and/or enhanced runoff, we argue that the Acarinina peak is not indicative of oligotrophic conditions. Instead, we postulate that (mainly low-trochospiral) Acarinina could have been better adapted to thrive in stressful surface water conditions than Morozovella, because it may have hosted different symbionts.

ACKNOWLEDGMENTS Christian Dupuis is thanked for providing samples of Dababiya. We are grateful to Dick Kroon, Jan Smit, and Martin van Breukelen for enabling stable isotope and carbonate measurements in Amsterdam. Saskia Kars and Karl-Heinz Baumann are thanked for help in taking SEM pictures in Amsterdam and in Bremen, respectively. Ralf Schiebel is kindly thanked for suggestions and discussions on ecology of extant planktic foraminifera. Eliana Fornaciari is thanked for nannofossil analysis of Gebel Qurtayssiat. E.G. warmly thanks Jochen Kuss for logistic support, and Claudia Agnini and Christian Scheibner for fruitful discussions on the PETM. Thoughtful suggestions from Estoquio Molina and an anonymous reviewer greatly improved the manuscript. This project was supported by the Deutsche Forschungsgemeinschaft (DFG) within the European Graduate College “Proxies in Earth History.” REFERENCES CITED Alegret, L., Ortiz, S., Arenillas, I., and Molina, E., 2005, Paleoenvironmental turnover across the Paleocene/Eocene boundary at the stratotype section in Dababiya (Egypt) based on benthic foraminifera: Terra Nova, v. 17, p. 526–536. Arenillas, I., and Molina, E., 1997, Análisis cuantitativo de los foraminíferos planctónicos del Paleoceno de Caravaca (Cordillera Bética): Bioestratigrafía y evolución de las asociaciones: Revista Española de Micropaleontología, v. 12, p. 207–232. Arenillas, I., Molina, E., and Schmitz, B., 1999, Planktic foraminiferal and δ13C isotopic changes across the Paleocene/Eocene boundary at Possagno (Italy): International Journal of Earth Sciences, v. 88, p. 352–364, doi: 10.1007/s005310050270. Aubry, M.-P., 1998, Early Paleogene calcareous nannoplankton evolution: A tale of climatic amelioration, in Aubry, M.-P., Lucas, S., and Berggren, W.A., eds., Late Paleocene–early Eocene climatic evolution and biotic events in the marine and terrestrial records: New York, Columbia University Press, p. 158–203. Aubry, M.-P., Berggren, W.A., Stott, L.D., and Sinha, A., 1996, The upper Paleocene–lower Eocene stratigraphic record and the Paleocene-Eocene boundary carbon excursion: Implications for geochronology, in Knox, R.W.O’B., Corfield, R., and Dunay, R.E., eds., Correlation of the early Paleogene in Northwest Europe: Geological Society [London] Special Publication 101, p. 353–380. Bains, S., Norris, R.D., Corfield, R.M., and Faul, K.L., 2000, Termination of global warmth at the Paleocene/Eocene boundary through productivity feedback: Nature, v. 407, p. 171–174, doi: 10.1038/35025035. Bé, A.W.H., 1977, An ecological, zoogeographic and taxonomic review of recent planktonic foraminifera, in Ramsay, A.T.S., ed., Oceanic micropaleontology, Volume 1: London, Academic Press, p. 1–100. Benjamini, C., and Sheva, B., 1992, The Paleocene–Eocene boundary in Israel—A candidate for the boundary stratotype: Neues Jahrbuch für Geologie und Paläontologie, Abhandlungen, v. 186, p. 49–61. Berger, W.H., 1970, Planktonic foraminifera: Selective solution and the lysocline: Marine Geology, v. 8, p. 111–138, doi: 10.1016/0025-3227 (70)90001-0. Berggren, W.A., and Norris, R.D., 1997, Biostratigraphy, phylogeny and systematics of Paleocene trochospiral planktic foraminifera: Micropaleontology, v. 43, p. 1–116, doi: 10.2307/1485988. Berggren, W.A., and Ouda, K., 2003, Upper Paleocene–lower Eocene planktonic foraminiferal biostratigraphy of the Dababiya section, Upper Nile Valley (Egypt), in Ouda, K., and Aubry, M.-P., eds., The upper Paleocene– lower Eocene of the upper Nile Valley, Part 1, Stratigraphy: Micropaleontology, v. 49, suppl. 1, p. 61–92, doi: 10.2113/49.Suppl_1.61. Berggren, W.A., Kent, D.V., Swisher, C.C., III, and Aubry, M.P., 1995, A revised Cenozoic geochronology and chronostratigraphy, in Berggren,

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The Geological Society of America Special Paper 424 2007

Calcareous nannofossil assemblages and their response to the Paleocene-Eocene Thermal Maximum event at different latitudes: ODP Site 690 and Tethyan sections Eugenia Angori Dipartimento di Scienze della Terra, Università degli Studi di Firenze, Via La Pira 4, 50121 Florence, Italy Gilen Bernaola Departamento de Estratigrafía y Paleontología, Facultad de Ciencia y Tecnología, Basque Country University, Apartado 644, E-48080 Bilbao, Spain Simonetta Monechi Dipartimento di Scienze della Terra, Università degli Studi di Firenze, Via La Pira 4, 50121 Florence, Italy

ABSTRACT A major change in calcareous nannofossil assemblages has been reported at the Paleocene-Eocene Thermal Maximum (PETM) on a global scale. To document the response of the nannoplankton communities below, within, and above the PETM, we studied in detail six successions, representing a wide range of environments and latitudes. Calcareous nannofossil response was different in discrete paleogeographic areas. Several classical Tethyan sections (Alamedilla, Caravaca, Zumaia [Spain], Contessa [Central Italy], and Wadi Nukhl [Egypt]), plus the high-latitude Ocean Drilling Program reference Site 690 (Weddell Sea) were re-investigated using high resolution calcareous nannofossil quantitative analyses. Five assemblage zones were identified: two before the onset of the Carbon Isotope Excursion (CIE) and three after it. Before the PETM, several changes were observed in both high and low latitudes that are characterized by well-defined increases of r-selected taxa (Biscutum and Prinsius). These changes probably were in response to an upwelling pulse that increased nutrients in surface waters. These events, which predate the geochemical and oceanic changes at the PETM, indicate that there were global events occurring before the actual CIE onset. At Site 690, the principal calcareous nannofossil change coincides with the onset of the CIE and is characterized by the rapid replacement of cold-water taxa by warm-water taxa. This change resulted from a sudden expansion of warmwater low-latitude assemblages into higher latitudes, probably due to an abrupt increase of surface-water temperatures. An increase in species richness here is due to the migration of several genera (i.e., Discoaster and Fasciculithus) south from warmer areas and to decreased dissolution. Moreover, an increase in abundance of Thoracosphaera spp. (calcareous dinoflagellate) below and within the CIE also indicates a stressed surface-water environment. Angori, E., Bernaola, G., and Monechi, S., 2007, Calcareous nannofossil assemblages and their response to the Paleocene-Eocene Thermal Maximum event at different latitudes: ODP Site 690 and Tethyan sections, in Monechi, S., Coccioni, R., and Rampino, M.R., eds., Large Ecosystem Perturbations: Causes and Consequences: Geological Society of America Special Paper 424, p. 69–85, doi: 10.1130/2007.2424(04). For permission to copy, contact [email protected]. ©2007 The Geological Society of America. All rights reserved.

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Angori et al. In the Tethyan sections, the response of the calcareous nannofossil assemblages to the PETM is more complex. As at the Southern Ocean Site 690, calcareous nannofossil fluctuations begin below the onset of the CIE and increase in frequency and amplitude at the benthic foraminifera extinction (BFE). At this level, calcareous nannofossil diversity and abundance abruptly decrease, and the Rhomboaster spp.– Discoaster araneus (R-D) association appears. The occurrence of the R-D association together with Thoracosphaera suggests that during the PETM there was a change to stressed ocean-surface conditions. Calcareous nannofossil recovery occurred later in the Tethys than at the southern high latitudes, where it occurred before the CIE recovery. Furthermore, the nannofloral assemblages after the δ13C recovery still indicate stressed conditions, suggesting that the plankton communities did not completely recover until later. Keywords: Paleocene/Eocene, calcareous nannofossils, paleoecology, Tethys, ODP 690.

INTRODUCTION The Paleocene-Eocene (P-E) boundary transition, characterized by a prominent negative excursion of the δ13C values (Carbon Isotope Excursion, CIE) in both marine and continental records (Stott et al., 1990; Kennett and Stott, 1991; Koch et al., 1992, 1995; Bralower et al., 1995; Thomas and Shackleton, 1996; Beerling and Jolley, 1998), has been the focus of significant interest by the scientific community in the last two decades. This interval is known as the PETM (Paleocene-Eocene Thermal Maximum) and is one of the most abrupt and transient extreme global climatic events to be documented in the ancient geologic record. According to recent studies, the PETM lasted for only ~220 k. y. (Röhl et al., 2000; Farley and Eltgroth, 2003) and was characterized by extremely warm climatic conditions, drastic changes in the global carbon budget, modifications in oceanic circulation and chemistry (possibly changes in the thermocline of the oceans), and reorganization of ecosystems. In fact, this event had an important biotic impact on both land and marine organisms, affecting groups as diverse as land mammals (Wood et al., 1941; Clyde and Gingerich, 1998; Gingerich, 2003), deep-water benthic foraminifers (Benthic Foraminifer Extinction, BFE) (Thomas, 1990, 1998; Thomas and Shackleton, 1996; Speijer et al., 1996), benthic foraminiferal assemblages of middle and outer neritic marine settings (Speijer, 1994; Speijer et al., 1996; Speijer and Schmitz, 2000), and shallow-marine larger foraminiferal fauna (Orue-Etxebarria et al., 2001; Pujalte et al., 2003). The impact of the PETM was significant in the photic zone, where various groups of planktic marine organisms suffered considerable turnovers. Dinoflagellates showed global dispersal and an increased abundance of species of the genus Apectodinium in coastal environments (Crouch et al., 2001). Planktic foraminifera experienced a radiation of shallow-dwelling tropical taxa (Kelly et al., 1996, 1998, 2005; Berggren and Ouda, 2003), and the calcareous nannoplankton communities underwent an important assemblage turnover that lasted for the duration of the event (Aubry, 2001; Bralower, 2002; Kahn and Aubry, 2004; OrueEtxebarria et al., 2004).

Although in the last few years several authors have published quantitative analyses to determine the response of calcareous nannofossils to the P-E event (Monechi et al., 2000a, 2000b; Bralower, 2002; Kahn and Aubry, 2004; Tremolada and Bralower, 2004; Orue-Etxebarria et al., 2004), the changes in abundance of calcareous nannofossils throughout the P-E transition are not yet well constrained. Moreover, a controversy still exists among authors whether the PETM was a period of increased or decreased productivity. Benthic foraminiferal accumulation rates and assemblages (which are dominated by species indicative of high productivity) suggest increased surface production and organic flux to the seafloor during the PETM (Thomas, 1998). This same conclusion can be drawn from the global dispersion of the Apectodinium spp. dinoflagellates (Bujak and Brinkhuis, 1998). However, geochemical productivity indicators, specifically increased Ba accumulation rates, have been interpreted to be the result of increased primary productivity during the PETM event (Bains et al., 2000), whereas Dickens (2001) believed this Ba increase to be a result of anaerobic oxidation of the large quantities of methane derived from hydrate dissociation. More recently, Stoll and Bains (2003), using the Sr/Ca ratios on coccolith carbonate, concluded that during the PETM phytoplankton productivity increased, an argument that has been questioned by Bralower (2004). Calcareous nannoplankton analyses also have yielded contradictory results. Bralower (2002), using calcareous nannofossil data from ODP Site 690 that show a replacement of r-mode specialists and taxa adapted to colder waters (i.e., Biscutum, Chiasmolithus) by warmer-water taxa (i.e., Discoaster, Fasciculithus), suggested a shift from colder, more productive surface waters to warmer more oligotrophic conditions during the PETM. This is at odds with evidence presented by Orue-Etxebarria et al. (2004) from the Zumaia section, where the major change in calcareous nannofossil assemblages at the PETM corresponds to the abrupt decrease in relative abundance of Zygrhablithus bijugatus, a species thought to indicate oligotrophic conditions (Aubry, 1998). According to those authors, the minimal abundance increases exhibited by Discoaster and Fasciculithus in

Calcareous nannofossil assemblages and their response to the PETM event connection with the PETM at Zumaia were probably controlled more by water temperature than by nutrient supply. To date, no one has been able to present any irrefutable conclusions about whether surface-water productivity increased or decreased during the PETM, and whether the plankton response to the PETM resulted mainly from temperature forcing or from changes in structure of the water column. With the aim of further documenting and more fully constraining the effects of the PETM in the calcareous plankton communities in a wider range of environments and latitudes, we reinvestigated and here summarize the response of calcareous nannoplankton assemblages below, within, and above the PETM in several classical Tethyan sections and at Southern Ocean ODP Site 690 (Fig. 1). We also compare the results obtained in these sections with those obtained by other authors from high-resolution calcareous nannofossil analyses in a wider range of latitudes (i.e., Bybell and Self-Trail, 1995, 1997; Bralower, 2002; Tremolada and Bralower, 2004; Kahn and Aubry, 2004, Gibbs et al., 2006). In particular, our goal is to characterize in detail the interval between the first drop in abundance of the genus Fasciculithus, 1–3 m below the CIE, and its extinction 5–16 m above the onset of the CIE. We will address the following topics: (1) features of the pre-boundary calcareous nannofossil assemblages; (2) changes or fluctuations in calcareous nannofossils before, during, and after the PETM, paying special attention to the quantitative determination of the changes undergone by the Rhomboaster spp.-Discoaster araneus association (R-D); and (3) recovery of the ecosystem. THE SECTIONS In this study, we have revisited and conducted higherresolution calcareous nannofossil quantitative analyses of six sections and sites: Alamedilla, Caravaca, Contessa, and Wadi Nukhl in the Tethyan realm, Zumaia in an intermediate position between the Tethys and the Atlantic Ocean, and finally the high southern latitude reference ODP Site 690 (Fig. 1).

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Most of the studied sections represent the P-E boundary transition at low latitudes from epicontinental (Alamedilla, Caravaca, and Wadi Nukhl) to open-ocean (Zumaia and Contessa) environments, whereas ODP Site 690 represents a high-latitude, deepwater marine environment. In the last few years, Site 690 has been the subject of micropaleontologic, cyclostratigraphic, and chemostratigraphic analyses (i.e., Bains et al., 1999; Röhl et al., 2000; Thomas and Bralower, 2001; Thomas, 2003; Kelly et al., 2005), and various paleo-productivity proxies have been tested to determine if the PETM was a period of weakened or strengthened surface-water productivity (Bralower, 2002, 2004; Stoll and Bains, 2003). For this study, a detailed analysis of calcareous nannofossils was carried out at this site from 2.5 m below to 4.5 m above the P-E boundary, which is located by definition at the onset of the CIE (Luterbacher et al., 2000). It is important to point out that at this site only the last common occurrence (LCO) of Fasciculithus has been reported, because the last occurrence (LO) of Fasciculithus occurs well above the studied interval (Pospichal and Wise, 1990). See details in the discussion paragraph. The P-E section at Alamedilla, located in the Subbetic Zone of the Betic Cordillera, is expanded and fairly continuous. Recently, it has been the focus of several studies that document important biotic, isotopic, and sedimentary changes that coincide with the P-E transition (i.e., Arenillas and Molina, 1996; Lu et al., 1996, 1998; Monechi et al., 2000a). From benthic foraminifer assemblages, a paleodepth of 1000–1200 m has been estimated (Arenillas and Molina, 1996; Lu et al., 1998). The studied interval ranges from 3 m below to 12 m above the onset of the CIE. The Caravaca section, geologically located in the same area, is a near-continuous Upper Cretaceous to middle Eocene marine sequence that has been studied in detail by several authors (i.e., von Hillebrandt, 1974; Romein, 1979; Canudo et al., 1995; Angori and Monechi, 1996). Its depositional depth of 600–1000 m corresponds to the inner slope. The studied interval ranges from 5 m below to 13 m above the P-E boundary.

B

Figure 1. (A) Paleoceanographic map (Lower Eocene; Scotese, 2001) showing the location of ODP Site 690 and Tethyan sections. (B) Detailed location of the studied Tethyan sections: Alamedilla, Caravaca, Contessa, Wadi Nukhl, and Zumaia.

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The P-E transition of the classic Contessa Road section in the Umbria-Marche Basin has been also studied in detail by various authors (Lowrie et al., 1982; Napoleone et al., 1983, Corfield et al., 1991; Galeotti et al., 2000; Coccioni et al., 2004). According to Kuhnt and Kaminski (1989), this section indicates a deep-basin setting at a depositional depth of 1500–2000 m. The studied interval ranges from 2.3 m below to 4 m above the P-E boundary. The outer-neritic Wadi Nukhl section in Egypt is located on a generally northwest-dipping paleoslope of an epicontinental basin. Benthic foraminifer assemblages indicate overall deposition at ~500 m (Speijer et al., 2000). To complete the detailed calcareous nannofossil biostratigraphic study carried out by Monechi et al. (2000b) in this section, we now present new quantitative analyses of this fossil group from 6 m below to 17 m above the onset of the CIE. The Zumaia section is one of the most expanded of known land-based deep-marine sections that contain the PETM, and, since the pioneer papers of Hillebrandt (1965) and Kapellos (1974), it has become a reference section for Paleogene bio-, chemo-, and cyclostratigraphic studies (i.e., Canudo and Molina, 1992; Schmitz et al., 1997; Baceta et al., 2000; Dinarès-Turell et al., 2002; OrueEtxebarria et al., 2004). The Paleogene succession was laid down in an offshore basin at an estimated water depth of ~1000 m (Pujalte et al., 1998 and references therein). The studied interval ranges from 3 m below to 12 m above the P-E boundary. METHODS In all sections and sites studied, the analysis of the calcareous nannofossil assemblage was carried out on centimeterspaced samples near the P-E boundary and at successively longer intervals above and below it. Smear slides of each sample were prepared from raw sediment, without centrifuging or shaking in order to avoid changes in the composition of the original assemblages. All the smear slides were analyzed under a Zeiss Axioplan II petrographic microscope at 1500× magnification (2000× for details and/or small taxa). For the quantitative analysis, at least 300 specimens per sample were counted in randomly selected fields of view. In most samples, Toweius pertusus and Coccolithus pelagicus exceeded 50% of the assemblage, thus obscuring the abundance changes of other, less prevalent species. Some of these taxa, although relatively scarce, have turned out to be essential (a key tool) for the interpretation of calcareous nannofossil association changes through the PETM. This is the case for species of the Rhomboaster-D. araneus association (Discoaster anartios, D. araneus, Rhomboaster spineus, and R. bramlettei s.a. and l.a.), a short-lived species association characterized by nannoliths with anomalous structure that lasted for the duration of the PETM (Aubry, 2001). To document the relative abundance changes of the representatives of Discoaster and Rhomboaster, we performed additional counts of those forms along a 31-mm track. For paleoecological purposes, and to reduce variables and enhance trends, we have included the discoasters other than D. anartios and D. araneus as Discoaster spp.

CALCAREOUS NANNOFOSSIL ASSEMBLAGE CHANGES Quantitative analyses revealed the presence of several different well-defined informal assemblage intervals that correspond to changes in calcareous nannofossils during the PETM time interval. Profound changes in nannofloras have been reported in previous studies at the onset of the PETM (Monechi et al., 2000a, 2000b; Bralower, 2002; Tremolada and Bralower, 2004; Kahn and Aubry, 2004; Orue-Etxebarria et al., 2004), where significant increases and decreases are primarily linked to several important genera of nannoliths and holococcoliths such as Discoaster, Rhomboaster, Fasciculithus, and Zygrhablithus. We have observed that the levels below, within, and above the PETM can be largely defined by the relative decrease and/or increase in abundance of Fasciculithus and Zygrhablithus, respectively (FasciculithusZygrhablithus abundance reversals). Five assemblage (intervals) zones were recognized: two below the onset of the CIE and three above it over a time interval of ~0.5 m.y. (Figs. 2–7). The upper Paleocene assemblages of the Tethyan sections are broadly comparable, though dominance of individual species can vary from one locality to another. These typical, low-latitude, upper Paleocene assemblages are named the Paleocene Assemblage Zone A, which is characterized by abundant C. pelagicus, T. pertusus, and representatives of the genera Fasciculithus (many species for this genus), Discoaster, and Sphenolithus. The calcareous nannofossil abundance and species richness are high in all the sections. At Zumaia, for example, the species richness ranges from 40 to 50 species, and the Shannon-Weaver diversity index (S-H) is always higher than 2.1. For a more detailed description of the assemblages, see Angori and Monechi (1996), Monechi et al. (2000b), and Orue-Etxebarria et al. (2004). At the high-latitude Site 690, the calcareous nannofossil association of the latest Paleocene is characterized by abundant Toweius and Chiasmolithus and by the presence of Biscutum, Sphenolithus, and Zygrhablithus bijugatus. The species richness, which does not vary markedly in all samples, is moderate and never exceeds 25. For a more detailed description of the assemblages, see Bralower (2002) and Angori and Monechi (2003). The boundary between the Assemblage Zone A and Paleocene Assemblage Zone B is marked by the beginning of a drop in Fasciculithus, in both abundance and species diversity, and by an increase in abundance of Zygrhablithus. In low latitudes these trends include a slight decrease in abundance of Discoaster and Sphenolithus, whereas in high southern latitudes these trends are not so distinct, and this interval is characterized only by an increase in Biscutum abundance. The increase in Biscutum has been recorded in several other sections (e.g., Site 690 by Bralower 2002; Site 401 by Tremolada and Bralower 2004). In Zone B of the Tethyan sections, the diversity generally continues to be high, but greater fluctuations occur in the number of species from sample to sample (from 36 to 51 species per sample). Both the Fasciculithus/Zygrhablithus reversal and the Biscutum peak of Zone B correspond to an interval of rapidly increasing δ18O values

Figure 2. Relative abundance of selected calcareous nannofossil taxa across the P-E transition in the Zumaia section, western Pyrenees. Percentages of Zygrhablithus, Fasciculithus, and Discoaster spp. are calculated from the non Toweius spp.+C. pelagicus fraction. Discoaster spp. refers to all Discoaster except D. anartios and D. araneus. Discoaster anartios, Discoaster araneus, Rhomboaster spineus, Rhomboaster bramlettei “long arms,” and Rhomboaster bramlettei “short arms” are represented in number of specimens in a 31-mm-long track. A, B, C, and D indicate assemblage zones (explanation in the text). Carbon isotopic record from Schmitz et al. (1997).

Figure 3. Relative abundance of selected calcareous nannofossil taxa across the P-E transition in the Alamedilla section (Betic Cordillera). Carbon isotopic record from Lu et al. (1998). See Figure 2 for further explanation.

Figure 5. Relative abundance of selected calcareous nannofossil taxa across the P-E transition in the Contessa Road section (Umbria-Marche Basin). Carbon isotopic record from Galeotti et al. (2000). See Figure 2 for further explanation.

Figure 4. Relative abundance of selected calcareous nannofossil taxa across the P-E transition in the Caravaca section (Betic Cordillera). Carbon isotopic record from Molina et al. (1994). See Figure 2 for further explanation.

Figure 6. Relative abundance of selected calcareous nannofossil taxa across the P-E transition in the Wadi Nukhl section (Egypt). Carbon isotopic record from Speijer et al. (2000). See Figure 2 for further explanation.

Figure 7. Relative abundance of selected calcareous nannofossil taxa across the P-E transition in the Site 690 (Maud Rice, Weddell Sea). Carbon isotopic record from Bains et al. (1999). See Figure 2 for further explanation.

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in surface-dwelling foraminifera (Bralower, 2002). Moreover, the base of this zone seems to coincide with the A– event, a slight decrease of the δ13C, which was defined by Zachos et al. (2005). The base of Eocene Assemblage Zone C coincides with the onset of the CIE, and it is marked by an increase of Discoaster and Fasciculithus and by a sharp decrease in the abundance of Toweius and Zygrhablithus (the latter especially in low latitudes). In the high southern latitudes, this assemblage change is abrupt and is characterized by the sudden replacement of Chiasmolithus and Biscutum by Fasciculithus and Discoaster, which, together with Toweius, are the dominant species throughout Zone C. In low latitudes, the appearance of Rhomboaster is also recorded at the onset of the CIE. Assemblage Zone C includes the entire CIE and, on the basis of the species recognized and their fluctuations in abundance, can be subdivided in two parts: a lower Subzone C-1 and an upper Subzone C-2. Subzone C-1 (Figs. 2–7) corresponds approximately to the interval between the onset of the CIE and the BFE (around the minimum carbon isotope values). It is discriminated at the base and top by two distinctive increases in abundance of Discoaster, mainly D. multiradiatus (along with rare D. nobilis, D. delicatus, and D. falcatus), and of Fasciculithus, mainly F. tympaniformis (along with rare F. cf. hayii and F. thomasii). In the Tethyan sections, Zygrhablithus shows a significant abundance decrease in an inverse correlation with Fasciculithus. This subzone (C-1) is quite distinct in the Zumaia, Alamedilla, and Contessa sections. In Caravaca, it is very thin with the CIE and BFE almost coincident (Molina et al., 1994; Pak and Miller 1992) because of sediment condensation or hiatus. At all Tethyan sections, the calcareous nannofossil richness decreases slightly throughout this subzone (C-1); in Zumaia the species richness varies from 40 to 30 and in the S-H index values from 2 to 1.3. This decrease is probably due to the presence of increased dissolution of specimens. At Site 690, this interval is delimited at its base by the abrupt abundance increase of Discoaster and at the top by the increase in abundance of Fasciculithus and by the occurrence of D. araneus. The superjacent Subzone C-2 corresponds to the interval between the BFE (around the minimum carbon isotope values) and the level where δ13C returns to pre-excursion values. It is distinguished by abundant Fasciculithus (which shows a marked decrease throughout this subzone) and by the absence or very low abundance of Zygrhablithus. At high latitudes (Sites 690B, 401), Zygrhablithus occurs in low abundance, whereas at low latitudes (Zumaia, W. Nukhl, Contessa, Alamedilla, Site 213, New Jersey) this genus is absent. In all analyzed low-latitude sections, Subzone C-2 has low calcareous nannofossil abundances and species diversities, but the most important feature that characterizes this interval is the presence of the D. araneus-Rhomboaster assemblage, an event restricted to the PETM of the Tethys-Atlantic province (Kahn and Aubry, 2004). According to Aubry and Sanfilippo (1999) and Cramer et al. (1999), the R-D assemblage consists of Rhomboaster spineus, R. calcitrapa, R. cuspis, Discoaster araneus, and D. anartios. For usage of the Rhomboaster-Tribrachiatus

linage, we follow the taxonomic remarks of Angori and Monechi (1996); therefore, the representatives of the genus Rhomboaster included in the R-D assemblage are Rhomboaster spineus, R. bramlettei “long arms,” and R. bramlettei “short arms.” Previous authors have concluded that the R-D assemblage is correlative with the earliest Eocene CIE and it has been proposed to constitute a stratigraphic proxy for this event (Kahn and Aubry, 2004). In our expanded Tethyan sections, however, the R-D assemblage appears a bit higher than the onset of the CIE, at the top of Subzone C-1, approximately at the level of the BFE around the minimum carbon isotope values, and disappears where the δ13C returns to pre-boundary values. However, very rare and scattered specimens of Rhomboaster have been identified at Zumaia (Orue-Etxebarria et al., 2004) and Wadi Nukhl sections (Figs. 2, 6) below the BFE. In those sections where D. anartios is present (Alamedilla, Caravaca, Contessa, Wadi Nukhl), it occurs together with D. araneus, but it is always rare and its occurrence is reduced at the bottom of the subzone (the interval where δ13C values show a decreasing trend). Discoaster araneus is regularly present throughout the entire subzone, but its maximum abundance occurs in the middle of the subzone (where δ13C values reach their minimum). As shown in Figures 2 through 7, R. bramlettei “long arms” has the same range as D. araneus, whereas D. spineus only appears sporadically. Rhomboaster bramlettei “short arms,” which has its FO before the other representatives of the R-D assemblage, shows its highest abundance in connection with the δ13C low values, and becomes less common with the extinction of the other representatives of the R-D assemblage. In fact, where δ13C returns to pre-excursion values, D. araneus, Rhomboaster spineus, and Rhomboaster bramlettei “long arms” disappear, and R. bramlettei “short arms” becomes less abundant. The only exception is the Wadi Nukhl section, where Rhomboaster increases again in abundance above the LO of D. araneus in the Zone D. At the high southern latitude Site 690, the presence of the R-D assemblage during the PETM is not so clear. Bralower (2002) reported the presence of D. araneus and D. anartios in Site 690 in the lower part of the δ13C excursion (between 170.66 and 170.4 meters below seafloor [mbsf]) and Rhomboaster was not found. The FO of Rhomboaster has been found higher up in the sequence (Pospichal and Wise, 1990; Angori and Monechi, 2003). New counts on a few smear slides of this interval confirm the presence of very rare specimens of D. cf. D. anartios at the CIE onset (170.66 mbsf) and show the occurrence of D. araneus from 170.46 up to 170.10 mbsf. Furthermore, it must be noted that D. araneus shows a great variability. Although they fit in with the original description, the forms found at Site 690 are quite different from those found in the Tethyan region (Fig. 8). Thoracosphaera is present with very low abundances within the δ13C excursion. In all of the Tethyan sections, Assemblage Zone D correlates with the return of δ13C to pre-excursion values and records important changes in calcareous nannofossil assemblages. Zygrhablithus reappears and increases considerably, becom-

Calcareous nannofossil assemblages and their response to the PETM event

Figure 8. Discoaster araneus morphologic variability between specimens from Tethyan sections and ODP Site 690. (1) Wadi Nukhl, 6.3. (2) Alamedilla, 14.5. (3) Caravaca, 33. (4) Zumaia, 20. (5) G. Qreya (Egypt), 25. (6) Site 690B-9-03, 30. Numbers represent the marks of the samples.

ing one of the most abundant taxa together with C. pelagicus, Toweius, and minor Sphenolithus, whereas Discoaster and Fasciculithus exhibit an important decrease in relative abundance. The R-D assemblage is no longer present, and only some representatives of R. bramlettei “short arms” can be found, except at Wadi Nukhl, where Rhomboaster increases in abundance and partly replaces Zygrhablithus, which is absent in this interval. The FO of R. bramlettei var. T (T. bramlettei sensu Aubry) occurs approximately at the base of this zone, where relative abundance increases of Neochiastozygus, Hornibrookina, and Braarudosphaera bigelowii (mainly in Zumaia) also occur. An exception is at Wadi Nukhl, where R. bramlettei var. T is absent throughout the studied interval. Together with all those changes, and in parallel with the increase in δ13C values, there is a strong increase in calcareous nannofossil abundance and species richness (e.g., in Zumaia the S-H index values reach 1.75 and 25–32 species per sample). In the high-latitude Sites 690 and 401, the recovery of Zygrhablithus and the relative abundance decrease of Discoaster and Fasciculithus occur before δ13C returns to preexcursion values. Sphenolithus shows an increase in abundance at the base of Zone D followed by a successive decrease, which is in agreement with Kelly et al. (2005). Rhomboaster bramlettei var. T has not been found at Site 690 up to 146.78 mbsf, 22 m above the recovery of Zygrhablithus (Angori and Monechi, 2003). Chiasmolithus and Toweius also characterize this interval; the former never again reaches the abundance it had before the PETM, whereas the latter achieves percentages greater than 70%. DISCUSSION In the last few years, several calcareous nannofossil paleoecologic analyses have been done in connection with the P-E transition in order to characterize the calcareous nannofossil turnover

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during the PETM (Monechi et al., 2000a, 2000b; Bralower, 2002; Kahn and Aubry, 2004; Tremolada and Bralower, 2004; OrueEtxebarria et al., 2004; Kelly et al., 2005). All of these works are based on the paleoecologic affinities attributed to the Paleocene and Eocene taxa in numerous paleobiogeographic studies of Paleogene nannofossils carried out over the last several decades (Haq and Lohmann, 1976; Haq et al., 1977; Wei and Wise, 1990; Aubry, 1992, 1998; Erba et al., 1992; Bralower, 2002). The paleoecologic affinities attributed to each taxon are disputed, and in some cases opposite relationships of some species with temperature and/or fertility have been proposed by different authors. For example, the holococcolithophorid Zygrhablithus, one of the most important taxa for defining calcareous nannofossil assemblage changes during the PETM, was interpreted as an oligotrophic species by Aubry (1998), but it reflects more eutrophic, cooler conditions according to Tremolada and Bralower (2004). Recent studies on living holococcolithophorids show their affinities for shallower environments, negligible turbulence, and normal nutrient conditions (Triantaphyllou et al., 2002). Moreover, it is well known that ecological preferences of individual taxa can change throughout time. There is, however, general agreement among authors that low total primary productivity and low nutrients correlate with high diversity of nannofossil assemblages. In fact, modern coccolithophorids are organisms that are adapted to an extremely low availability of nutrients, and today their maximum diversity is observed in tropical oligotrophic environments, where the vertical distribution of the species indicates stratification of the photic zone (Kilham and Soltau Kilham, 1980). The five assemblage zones recognized in the P-E transition in our sections represent different stages in calcareous-nannofossil evolution during the PETM. In the Tethyan area, the late Paleocene calcareous Assemblage Zone A is characterized by high and relatively stable calcareous nannofossil diversity and abundance values, indicative of generally stable oceanic conditions, a stratified water column, and low nutrient content. In modern oceans, these conditions can be found in oligotrophic areas such as the central gyres, which are characterized by a stable water column, a deep thermocline, and diverse phytoplankton communities dominated by specialized species. Immediately before the PETM, in Assemblage Zone B, the stability of calcareous nannofossil assemblages ends. In low latitudes, Fasciculithus, Discoaster, and Sphenolithus begin to decrease, and Zygrhablithus bijugatus exhibits different behavior depending on setting. It is absent at Wadi Nukhl (Fig. 6), present in very low abundance at Alamedilla and Caravaca (Figs. 3 and 4), and abundant in deeper sections as well as at high southern latitudes (Site 690). An increase in abundance of Biscutum and Prinsius bisulcus has been found in this interval. These peaks correspond to an interval of rapidly increasing δ18O values in surface-dwelling foraminifers (Kennett and Stott, 1991), and were probably produced by a pulse of increased upwelling that elevated nutrient supply in surface waters. These events, recorded both in the Tethys and at Site 690, were a prelude to

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the geochemical and oceanic changes at the PETM and indicate that changes began before the CIE onset (i.e., the massive release of methane hydrates). In high latitudes, an abrupt and rapid change in the calcareous nannofossil assemblages occurred in connection with the onset of the CIE. Chiasmolithus and Biscutum were suddenly replaced by Discoaster and Fasciculithus, which showed abrupt increases in their relative abundances. Although this change began gently a bit earlier, it became abrupt when the δ13C values suddenly decreased, indicating a shift from cold-water to warmer-water taxa (Bralower, 2002). The assemblage is dominated by Discoaster and Fasciculithus throughout Zone C. Species diversity increases as taxa migrated south from warmer areas (D. falcatus, F. thomasii, F. cf. F. richardii, F. cf. F. hayii) and with the first occurrence of taxa (D. cf. D. anartios, D. araneus, D. salisburgensis, and Discoaster with a “prominent stem” [e.g., D. megastypus and D. multiradiatus]) characterized by analogous morphological features (large, compact, robust). In the Tethyan realm, changes recorded in the calcareous nannofossil assemblage at the PETM are more complex (Subzones C-1 and C-2). When the δ13C values abruptly decreased (CIE onset), a period of wide fluctuations in calcareous nannofossil assemblages began that lasted until the BFE (Subzone C-1). The beginning of this phase of oscillation in these assemblages in the Tethys occurs in parallel to the abrupt turnover of calcareous nannofossils in the high latitudes, which can be explained as a sudden expansion of low-latitude warm-water assemblages to high latitudes, probably as a result of an abrupt increase in the surface-water temperatures. The fluctuations in the assemblages recorded in the Tethyan sections could reflect different pulses of that expansion, as watercolumn conditions deteriorated in the upper Paleocene. In the Tethyan realm, the major calcareous nannofossil turnover recorded in the PETM is in connection with the BFE (Subzone C-2), where both diversity and abundance values decrease. These changes seem to be mainly due to a more or less intense episode of dissolution. From this level, and through the entire Subzone C-2, low diversity (essentially at Zumaia) and minimum calcareous nannofossil abundance is maintained. The low diversity and abundance registered in the entire interval, in several sections, could be the result of dissolution (the presence of a more or less expanded lysocline or calcite compensation depth). In contrast, at Site 690, species richness increases sharply. This can be explained by the migration of warm-water taxa (Discoaster and Fasciculithus) to high latitudes and by the reduced dissolution recorded at high latitude. Furthermore, the presence within Subzone C-2 of the r-selected taxa Cyclagelosphaera, Hornibrookina, and Neochiastozygus (Gibbs et al., 2006), suggests higher temperatures associated with higher nutrient supply (a mixed upper water column, shallow thermocline). Furthermore, within this interval, Discoaster and Fasciculithus, considered to be warm-water taxa and K-specialists (Bralower, 2002; Aubry, 1998), show an opposite trend. Orue-Etxebarria et al. (2004) suggested that the distribution of Discoaster is linked to surface-water temperature. Furthermore, as observed by Fenner (1986), during

the Eocene–Oligocene at the high southern latitudes (Site 511), a distinct temperature change occurred, but the nutrient availability at those latitudes did not vary. In fact, only warm-water-loving diatom species disappear. In this time interval, Discoaster also disappears in high latitudes; it therefore seems that the nutrients were not the limiting factor for discoasters, and that their disappearance was primarily linked to temperature changes. The most important change in the calcareous nannofossil assemblage during Subzone C-2 at low latitudes is the appearance of the R-D assemblage, a short-lived species association characterized by nannoliths with anomalous structure that according to Aubry (2001) occurred only in a crescent-shaped province that extends from the Caribbean through the Tethyan region to the westernmost Indian Ocean. The R-D assemblage, which suddenly appears at the BFE, is more abundant at the base of Subzone C-2, and it becomes less abundant as the δ13C returns to pre-boundary values. The origin and paleoecologic affinities of most of the representatives of the R-D assemblage are unknown, especially those of the Rhomboaster-Tribrachiatus lineage. Romein (1979), on the basis of similarities in construction, suggested that Micula decussata was the ancestral form of this lineage, which thereby implies that this taxon survived the K-P boundary mass extinction. Judging from the shape and morphological features of some forms in the R-D assemblage, we suggest that at least D. araneus, D. anartios, R. bramlettei “long arms,” and R. spineus, as well as discoasters with prominent stems, should be interpreted as malformed forms produced by the increase in CO2 concentration. The presence/abundance of these taxa could indicate regions with more or less CO2 concentration. The occurrence of Thoracosphaera, the new short-lived species association (which is composed probably of ecophenotypes), and its higher abundance values in this interval (Subzone C-2) suggest that during the PETM a change to stressed ocean-surface conditions occurred. At high-latitude successions, such as Site 690, the presence of the R-D assemblage is not normally noticeable, and the most important event in connection with the PETM is the replacement of Biscutum and Chiasmolithus by Discoaster and Fasciculithus, an assemblage change that, according to Bralower (2002), suggests a shift from a colder, more productive water column to a warmer, more stable stratified water column with increased oligotrophy. However, Stoll and Bains (2003), using Sr/Ca ratios of coccolith carbonate as a proxy to estimate phytoplankton productivity during the PETM, concluded that during this interval productivity increased at Site 690. This conclusion concurs with benthic foraminifer assemblage data (Thomas and Shackleton, 1996) and the observed increases in the Ba accumulation rates (Bains et al., 2000) at this site. To shed light on this controversy, Bralower (2004) recently calculated carbonate accumulation rates (CAR) across the PETM at Site 690. He found that at the onset of the CIE there were significant decreases in CAR that persisted for the first 50 k.y. of the event, the same interval characterized by increasing Sr/Ca values. He argued that the decrease in CAR is consistent with

Calcareous nannofossil assemblages and their response to the PETM event a decrease in bulk nannoplankton calcification rates, which in turn would seem to be more consistent with a weakened biological pump than a strengthened one as argued by Stoll and Bains (2003). Recent studies (Riebesell et al., 2000; Rost and Riebesell, 2004) indicate that the productivity and distribution of coccolithophores are very sensitive to the rise of CO2 in the oceans. The response of this group suggests a decrease in biocalcification and an increase in photosynthesis with a consequently lower PIC/POC ratio (Particulate Inorganic Carbon/Particulate Organic Carbon ratio). These changes could explain in part the decline of carbonate content (to a certain extent related to the shoaling of the lysocline) and the high amount of organic material in the deep waters (Thomas, 1998; Thomas et al., 2000). In the Tethyan sections, when δ13C returns to pre-boundary values (lowermost part of Zone D), the calcareous nannofossil association experiences another major turnover. The R-D assemblage is no longer present, Fasciculithus and Discoaster relative abundance decreases, and Zygrhablithus abruptly reappears. The calcareous nannofossil diversity and abundance increase but do not reach pre-boundary values. The S-H index (e.g., at Zumaia) never exceeds 1.7, and only 25–32 species per sample are recorded. These changes indicate an important turnover in the ocean water-column structure, probably from unstratified and productive water to progressively stratified, clearer, and less productive waters. This scenario could also explain the lower foraminifer shell content observed by Kelly et al. (2005) at Site 690. They suggested that the scarcity of foraminifera is due to a dilution effect caused by enhanced production/preservation of coccolithophorids. They also suggested that increased coccolithophorid carbonate production could explain a non-placolith (Discoaster, Fasciculithus, Sphenolithus) decline. Nevertheless, at Site 690 our data show that while the previous taxa decline, Zygrhablithus bijugathus (a non-placolith) markedly increases to reach abundances of around 50% of the assemblage (Toweius spp. removed). Thus, the decline of Discoaster, Fasciculithus, and Sphenolithus could also be linked to other features (e.g., temperature). The relative low diversity and abundance values of placoliths (mainly Toweius and C. pelagicus) in Zone D indicate continuing mesotrophic conditions. In fact, in some Tethyan sections, at the base of Zone D, there is a relative abundance increase of Hornibrookina, Neochiastozygus, and Braarudosphaera bigelowi (the latter in Zumaia), taxa considered by many authors to be indicative of eutrophic to mesotrophic conditions (Aubry 1998; Bralower 2002; Cunha and Shimabukuro, 1997; Gibbs et al., 2006). Abundance peaks of these taxa have also been found in eutrophic conditions in the lowermost Paleocene of various Cretaceous-Paleogene (K-P) low-latitude successions. The B. bigelowii abundance peak registered at those K-P boundary sections always follows the abundance peak of Thoracosphaera and occurs in connection with the disappearance of the “Strangelove” ocean conditions related to the extinction event. The absence of B. bigelowi at the PETM (Zone C), interpreted as a high-productivity interval, could be related to the presence of special or Strangelove conditions in the oceans at this time,

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as was suggested earlier by the presence of the R-D assemblage and by the increase in Thoracosphaera. The calcareous nannofossil turnover in Zone D has not been registered at the same level in the Tethys and in the high latitudes; in the latter it occurs earlier, before the recovery of the δ13C to pre-boundary conditions. The response of the calcareous nannofossil assemblages to the PETM is diverse in different paleogeographic areas. It seems to have been more complex and more intense and to have lasted longer in the Tethys than in the higher latitudes, where the R-D assemblage is not so noticeable and the recovery of the assemblage occurred before the recovery of δ13C to pre-boundary values. CONCLUSIONS Associated with the PETM is an important worldwide turnover in calcareous nannofossils. The detailed analysis carried out in this study, which documents five assemblage zones, two below the onset of the CIE and three above it, revealed that this response is different in discrete paleogeographic areas (Fig. 9). Before the PETM, several changes occurred in both high and low latitudes. These well-defined increases of r-selected taxa (Biscutum and Prinsius) probably were produced by a pulse of increased upwelling that elevated nutrient supply in surface waters. These events predate the geochemical and oceanic changes at the PETM, indicating forces at work before the onset of the CIE. In the high southern latitude Site 690, the calcareous nannofossil turnover is mainly characterized by the replacement of cold-water taxa by warm-water taxa. This replacement is probably due to an abrupt increase in surface-water temperatures with a resulting expansion of warm-water, low-latitude assemblages into higher latitudes. The increase in species richness is due both to the migration south of several taxa (Discoaster and Fasciculithus) from warmer areas and to reduced dissolution at this high-latitude site. The worldwide decline in CaCO3 content at the PETM probably is linked to a decrease in biocalcification and increased coccolithophorid photosynthesis, which were caused by an increase in the CO2 content in the water (Rost and Riebesell, 2004) and an increase in dissolution associated with the transient shoaling of the lysocline. In the Tethys, the response of the calcareous nannofossil association to the PETM is more complex. In high southern latitudes, the turnover began at the onset of the CIE but became much more abrupt at the benthic foraminifera extinction (BFE). At this level, calcareous nannofossil diversity and abundance abruptly decreased, and the R-D assemblage and Thoracosphaera appeared, suggesting that during the PETM ocean-surface conditions became stressed. In the Tethys, the calcareous nannofossil initial recovery occurred somewhat later than at high latitudes. However, both the Tethyan and high-latitude assemblages found after the δ13C recovery are indicative of mesotrophic conditions, suggesting that the plankton communities were not yet completely recovered.

Figure 9. Calcareous nannofossil assemblages across the P-E transition: a correlation between Site 690 and Tethyan sections. The light gray band indicates the core of the negative carbon isotope excursion (CIE) interval. A, B, C, and D indicate assemblage zones (explanation in the text).

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MANUSCRIPT ACCEPTED BY THE SOCIETY 20 NOVEMBER 2006

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The Geological Society of America Special Paper 424 2007

A review of calcareous nannofossil changes during the early Aptian Oceanic Anoxic Event 1a and the Paleocene-Eocene Thermal Maximum: The influence of fertility, temperature, and pCO2 Fabrizio Tremolada* RPS Energy, RPS Group PLC, Woking, Surrey GU21 3LG, UK Elisabetta Erba Dipartimento di Scienze della Terra “Ardito Desio,” Università di Milano, 20133 Milano, Italy Timothy J. Bralower Department of Geosciences, Pennsylvania State University, University Park, Pennsylvania 16802, USA ABSTRACT The comparison between calcareous nannofossils during the early Aptian Oceanic Anoxic Event 1a (OAE1a) and the Paleocene-Eocene Thermal Maximum (PETM) suggests different nannofloral reactions to extreme greenhouse conditions. Both events were likely characterized by major changes in nutrient concentrations, temperature, and pCO2 levels. OAE1a corresponds to an increase in opportunistic taxa associated with eutrophic surface-water conditions. Eutrophy also resulted in the demise of an oligotrophic group, the nannoconids. Nannofloral assemblages of the PETM interval suggest nutrient-depleted surface waters at open-ocean sites including those at high and low latitudes. However, the upper part of the PETM shows a return to mesotrophic conditions documented by the increase in abundance of mesotrophic taxa. PETM records from shelf sites are characterized by an increase in nannofossil taxa indicative of mesotrophic conditions, suggesting an increase in productivity. Fluctuations in primary productivity affected composition and abundance of calcareous nannofossil assemblages during both events. Whereas fertility increased in the global ocean during OAE1a, mesotrophic conditions mostly characterized proximal settings during the PETM. Nannofloral changes could have been partially triggered by the warming, but the influence of high pCO2 levels is not evident. Reductions in nannofossil calcification and paleofluxes are associated with the OAE1a, but the role of pCO2 variations in nannofloral calcification during the PETM is not obvious. In both events, variations in lysocline/CCD depth and enhanced dissolution and/or diagenesis strongly affected nannofossil assemblages in some locations, but the overall nannofloral changes reveal a primary paleoecological and paleoceanographic signal. Keywords: Oceanic Anoxic Event 1a; Paleocene-Eocene Thermal Maximum; calcareous nannofossils; carbon dioxide; nutrient availability; temperature. *[email protected] Tremolada, F., Erba, E., and Bralower, T.J., 2007, A review of calcareous nannofossil changes during the early Aptian Oceanic Anoxic Event 1a and the PaleoceneEocene Thermal Maximum: The influence of fertility, temperature, and pCO2 , in Monechi, S., Coccioni, R., and Rampino, M.R., eds., Large Ecosystem Perturbations: Causes and Consequences: Geological Society of America Special Paper 424, p. 87–96, doi: 10.1130/2007.2424(05). For permission to copy, contact [email protected]. ©2007 The Geological Society of America. All rights reserved.

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Both events are also characterized by a major perturbation of the global C cycle. OAE1a is associated with a large positive carbon isotope excursion in both carbonate and organic matter (e.g., Jenkyns, 1980; Weissert, 1989; Arthur et al., 1990; Bralower et al., 1994, 1999; Menegatti et al., 1998; Moullade et al., 1998; Weissert et al., 1998; Erba et al., 1999; Renard et al., 2005). However, the lower part of the OAE1a shows a significant decrease to more negative δ13C values in both marine (e.g., Jenkyns, 1995; Menegatti et al., 1998; Moullade et al., 1998; Bralower et al., 1999; Erba et al., 1999; Bellanca et al., 2002) and terrestrial organic matter (Gröcke et al., 1999; Jahren et al., 2001). The PETM coincides with a prominent negative carbon isotopic excursion (CIE), which has been identified in sections from all ocean basins (e.g., Kennett and Stott, 1991; Thomas and Shackleton, 1996; Kaiho et al. 1996, Dickens et al. 1997; Bralower et al., 1997; Bains et al., 1999; Katz et al., 1999; Zachos et al., 2003) and from the terrestrial realm (e.g., Koch et al., 1992; 1995). The magnitude and rate

The early Aptian Oceanic Anoxic Event 1a (OAE1a) ca. 120 Ma (e.g., Larson and Erba, 1999) and the Paleocene-Eocene Thermal Maximum (PETM) ca. 55 Ma (e.g., Röhl et al., 2000) represent two of the most significant global climate and environmental perturbations of the last 140 m.y. (Figs. 1 and 2). OAE1a and the PETM are characterized by greenhouse climates that were likely caused by increased pCO2 levels. The causes of greenhouse conditions were possibly related to the formation of two large igneous provinces. The early Aptian interval was characterized by the emplacement of the giant Ontong Java and Manihiki Plateaus and the formation of the Nova Canton Trough in the Pacific Ocean (e.g., Larson, 1991a, 1991b; Larson and Erba, 1999). The PETM interval may overlap with the formation of the North Atlantic Igneous Province (e.g., Owen and Rea, 1985; Eldholm and Thomas, 1993; Svensen et al., 2004).

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of the CIE and the negative excursion associated with the OAE1a have been interpreted as the result of a massive injection of methane from the dissociation of clathrates (e.g., Dickens et al., 1995, 1997; Bralower et al., 1997; Opdyke et al., 1999; Jahren et al., 2001; Bellanca et al., 2002; Jenkyns, 2003). The goal of this study is to review major changes in calcareous nannofossil assemblages triggered by variations in nutrient availability, temperature, and possibly carbon dioxide concentrations across the OAE1a and the PETM. Calcareous nannofossils are an important source of calcite to pelagic carbonates, and fluctuations in abundance and composition of nannofloral assemblages may therefore affect carbonate accumulation. Major changes in abundance and composition of living nannoplankton communities are generally triggered by variations in primary productivity and temperature (e.g., Honjo and Okada, 1974; Winter et al., 1994; Brand, 1994; Roth, 1994). Similarly, major changes in nannofossil assemblages can be used to reconstruct variations in the thermal and trophic regime of ancient oceans. The influence of increasing carbon dioxide concentrations on abundance and composition of nannofloral communities is still unclear. Changes in calcareous nannofossil assemblages during both events have been documented in great detail (e.g., Erba, 1994; Larson and Erba, 1999; Erba and Tremolada, 2004; Bralower et al., 1994; Bralower, 2002; Tremolada and Bralower, 2004; Gibbs et al., 2006). Here we discuss similarities and differences in the response of nannoplankton communities to oceanographic and

environmental changes during OAE1a and PETM, and speculate on the role of temperature, primary productivity, and pCO2 levels as causal factors. THE OCEANIC ANOXIC EVENT 1A Major changes in abundance and composition of calcareous nannofloras characterize the Barremian-Aptian boundary interval (Fig. 3). In particular, this time interval records a global speciation episode in calcareous nannofossil communities that began 1.5 m.y. before the early Aptian OAE1a and continued through the late Aptian (e.g., Roth, 1989; Bralower et al., 1994; Erba, 1994, 2004). Causes for accelerations in evolutionary rates such as originations are still under debate (Bown et al., 2004; Erba, 2006), but it is difficult to discriminate among environmental triggers. The most marked fluctuations occur in the nannoconids, a group that dominates Early Cretaceous nannofossil assemblages, especially at low latitudes, and are thought to be oligotrophic (e.g., Busson and Nöel, 1991; Coccioni et al., 1992; Mutterlose, 1996). Several studies suggest that nannoconids are deep-dwelling forms and their variations in abundance are controlled by fluctuations in nutricline depth (e.g., Erba, 1994; Herrle, 2003; Watkins et al., 2005). However, high abundances of nannoconids have also been interpreted as indicative of increasing primary productivity (Scarparo Cunha and Koutsoukos, 1998; Street and Bown, 2000). Narrow-canal nannoconids dominate Barremian assemblages, whereas wide-canal

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Figure 3. Bulk δ13C at Cismon Drillcore, northeastern Italy (Erba et al., 1999) and major calcareous nannofossil changes across the OAE1a at low latitudes. Nannofossil data are a synthesis from Erba (1994), Bralower et al. (1994), Aguado et al. (1999), Larson and Erba (1999), Channell et al. (2000), and Erba and Tremolada (2004).

nannoconids increase in abundance in the uppermost Barremian and overwhelm the narrow-canal forms in the Aptian (e.g., Channell et al., 2000; Erba and Tremolada, 2004). Both groups experienced a major “crisis” some 40 k.y. prior to the OAE1a (e.g., Erba, 1994). Heavily calcified nannoconids are virtually absent in the OAE1a black shales, but coccoliths (especially opportunistic taxa) and other nannoliths such as normal- and over-sized Assipetra and Rucinolithus thrived (e.g., Erba, 1994; Aguado et al., 1999; Luciani et al., 2001; Bellanca et al., 2002; Erba and Tremolada, 2004). Assipetra and Rucinolithus are particularly abundant at the base of the OAE1a, coinciding with the negative spike in δ13C. Because these nannoliths correlate inversely with the nannoconids, Tremolada and Erba (2002) postulated that Assipetra and Rucinolithus had eutrophic affinities, although their ultrastructure is completely different from that of other high-productivity nannofossils. The return of nannoconids is observed above OAE1a, but this group never regained its former dominance (Erba, 1994). The nannoconid crisis and the increase in abundance of opportunistic calcareous nannofossil taxa suggest increasing pri-

mary productivity during OAE1a. This interpretation is supported by changes in radiolarian and planktonic foraminiferal communities (Premoli Silva et al., 1999) and in organic-walled phytoplankton assemblages (Hochuli et al., 1999). Biomarker data led Kuypers et al. (2004) and Dumitrescu and Brassell (2005) to suggest that cyanobacterial N2 fixation was the main source for nutrient N during OAE1a. The increasing fertility of surface waters during the OAE1a may have been the result of the formation of the Ontong Java and Manihiki Plateaus. This volcanism introduced huge quantities of carbon dioxide into the atmosphere, causing warm and humid (greenhouse) conditions (e.g., Weissert, 1989; Weissert et al., 1998; Larson and Erba, 1999), possibly further fueled by the methane dissociation and its oxidation (Opdyke et al., 1999; Jahren et al., 2001; Bellanca et al., 2002; Jenkyns, 2003; Price, 2003). Excessive carbon dioxide levels led to intensified runoff and continental weathering. Increased fluxes of nutrients may have caused the increase in fertility in proximal settings that resulted in an expanded oxygen minimum zone and dysoxic/anoxic bottom waters, leading to enhanced organic car-

Review of calcareous nannofossil changes bon burial (e.g., Weissert, 1989; Bralower et al., 1994; Erba et al., 1999). Furthermore, hydrothermal plumes added large amounts of toxic and biolimiting metals to marine ecosystems. High quantities of elements such as Fe, V, Ti, Zn, Mo, Mn, Ag, and Ni may have induced a large increase in primary productivity (e.g., Coale et al., 1996; Behrenfeld and Kolber, 1999). Significant concentrations of trace metals are recorded just below and within OAE1a black shales (R. Duncan, 1999, personal communication; Larson and Erba, 1999; Erba, 2004). The occurrence of biomarkers associated with N-fixing cyanobacteria in OAE1a black shales in both the Tethys and Pacific Oceans indicates nutrient cycling under increased iron availability (Kuypers et al., 2004; Dumitrescu and Brassell, 2005). THE PALEOCENE-EOCENE THERMAL MAXIMUM

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to be mesotrophic, such as Hornibrookina, Coronocyclus, Campylosphaera, Neochiastozygus, and Toweius (Gibbs et al., 2006). Mesotrophic nannofloral communities are commonly associated with high abundances of the dinoflagellate genus Apectodinium, which is interpreted to be adapted to warm and more nutrient-rich waters (Crouch et al., 2001; Crouch and Brinkhuis, 2005). Conversely, heavily calcified taxa such as Discoaster and Fasciculithus decrease abruptly. The genus Discoaster was an abundant component of Cenozoic low-latitude environments and probably adapted to warm-water conditions (e.g., Edwards, 1968; Bukry, 1973; Haq and Lohmann, 1976; Wei and Wise, 1990; Firth and Wise, 1992). Variations in relative abundance of Discoaster in tropical sites may suggest that this taxon thrived in oligotrophic conditions (Backman, 1986; Aubry, 1992). Also, Aubry (1992) suggested that this genus was a deep-dwelling form that expanded its range to high latitudes during warming events. The taxon Fasciculithus probably had ecological affinities similar to those of Discoaster. In particular, Haq and Lohmann (1976) documented an inverse correlation between Fasciculithus and Prinsius martinii, a taxon that flourished in eutrophic environments at high latitudes. Open-ocean settings show marked differences at low- and high-latitude settings. The CIE onset at low latitudes is commonly characterized by high abundances of Discoaster and

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Calcareous nannofossil assemblages show profound changes in abundance and composition across the Paleocene-Eocene transition (e.g., Bralower, 2002; Tremolada and Bralower, 2004; Youssef and Mutterlose, 2004; Gibbs et al., 2006). Various assemblage shifts are associated with different latitudinal and environmental settings (Fig. 4). Shelf areas are generally characterized by increasing abundances of small-sized taxa that are presumed

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Figure 4. Bulk δ13C at ODP Site 690, Weddell Sea (Bains et al., 1999), and major calcareous nannofossil changes across the PETM in proximal settings and open-ocean sites at both low and high latitudes. Nannofossil data are a synthesis from Bralower (2002), Tremolada and Bralower (2004), Orue-Extebarria et al. (2004), Raffi et al. (2005), and Gibbs et al. (2006).

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Fasciculithus, suggesting warm and oligotrophic conditions (e.g., Orue-Extebarria et al., 2004; Gibbs et al., 2006). High abundances of Fasciculithus were observed in northern Italy (Agnini et al., 2006) and in the eastern Pacific Ocean (Raffi et al., 2005). However, Monechi et al. (2000) documented a decrease in abundance of Fasciculithus just below the CIE at Alamedilla, Spain. The paucity of mesotrophic taxa, coupled with the abundance of taxa indicative of highly stressed environments such as calcareous dinoflagellates Biantholithus and Braarudosphaera, supports the interpretation of warm and nutrient-depleted surface waters (Gibbs et al., 2006). Similarly, the lower part of the PETM at high latitudes is characterized by the increase in abundance of warm and oligotrophic taxa such as Discoaster and Fasciculithus (Bralower, 2002; Tremolada and Bralower, 2004). In the upper part of the event, sharp increases in Toweius, Zygrhablithus, and Sphenolithus correspond to abrupt decreases in Discoaster and Fasciculithus. This major switch in nannofloral assemblages was likely caused by changes in thermal and nutrient structure of the water column during the PETM (Bralower, 2002; Tremolada and Bralower, 2004), leading to an increase in primary productivity. The formation of the North Atlantic Igneous Province possibly heightened greenhouse conditions during the PaleoceneEocene transition (Eldholm and Thomas, 1993) and possibly led to the release of methane hydrates (Svensen et al., 2004). The available geochemical, sedimentological, and micropaleontological information (see Bralower, 2002 and Zachos et al., 2003 for a synthesis) suggests that the PETM conditions affected shelf and open-ocean settings. In shelf settings, intensified runoff and chemical weathering (Robert and Kennett, 1994) introduced high concentrations of nutrients that led to higher productivity and, at some locations, combined with low oxygen concentrations, resulting in high total organic carbon contents (e.g., Bolle et al., 2000). Open-ocean sites do not show increasing TOC concentrations, although evidence of dysoxic seafloor conditions such as fine-laminated sediments are recorded at ODP Sites 999 and 1257 (Bralower et al., 1997; Erbacher et al., 2004). Unchanged TOC concentrations may suggest that organic fluxes from the photic zone were negligible, thus supporting the hypothesis of warm, oligotrophic and stratified surface waters. In addition, no increases in biolimiting and/or toxic metals are documented through the PETM (Thomas and Bralower, 2005). Nevertheless, other proxies could be indicative of eutrophic conditions in open-ocean sites. Benthic foraminiferal assemblages are thought to be indicative of anoxia and/or high food supply at the seafloor (Thomas and Shackleton, 1996; Speijer and Schmitz, 1998). Also, according to some authors (Bains et al., 2000; Stoll and Bains, 2003), the increase in Ba accumulation rates and a positive shift in the Sr/Ca ratio within the PETM in Hole 690B point to an increase in surface-water fertility. However, the increase in Ba concentrations may have resulted from the introduction of non-biogenic Ba ions stored along with methane hydrates (Dickens et al., 2003), and the use of the Sr/Ca ratio as paleoproductivity indicator is still under debate (Stoll and Bains, 2003; Bralower et al., 2004; Stoll, 2004).

PRIMARY PRODUCTIVITY, TEMPERATURE, AND CARBON DIOXIDE The OAE1a and PETM represent two major climatic and trophic perturbations that globally affected the earth. Calcareous nannofossils responded very differently to these two events. The Barremian-Aptian boundary records several new occurrences of calcareous nannofossils but no evidence of extinction in cosmopolitan taxa (Erba, 2004). The extinction of endemic taxa such as Nannoconus borealis and N. abundans predates the deposition of OAE1a black shales (e.g., Deres and Achéritéguy, 1980) and might reflect environmental stress in biogeographically restricted areas. The study of PETM nannofloral assemblages has revealed the entry of new short-range taxa such as the genus Rhomboaster and the species Discoaster araneus and D. anartios, considered “excursion taxa” (e.g., Monechi et al., 2000; Kahn and Aubry, 2004; Orue-Extebarria et al., 2004; Youssef and Mutterlose, 2004), which possibly filled short-lived ecological niches. In addition, Bralower (2002) and Agnini et al. (2006) documented the elimination of the genus Fasciculithus, a very abundant component of Paleocene nannofloral assemblages, likely outcompeted by the holococcolith Zygrhablithus to occupy the same paleoecological niche. Quantitative analyses of calcareous nannofossils suggest major changes in nutrient availability through the OAE1a and the PETM. Abundance and composition of nannofloral assemblages seem to have been affected predominantly by trophic conditions through both events. OAE1a nannofloras imply that all oceanic environments were characterized by enhanced primary productivity (Erba, 2004). Increasing quantities of nutrients from hydrothermal plumes, runoff, and weathering possibly stimulated the primary productivity in both shelf and oceanic sites (Larson and Erba, 1999; Leckie et al., 2002; Erba, 2004). Increasing nutrient concentrations could have led to a rising of the nutricline, which resulted in the “nannoconid crisis” (Erba, 1994; 2004). Conversely, during the PETM, abundance and composition of calcareous nannofossil communities fluctuated differently in shelf and in open-ocean sites. Oligotrophic taxa such as Discoaster and Fasciculithus thrived in open-ocean sites, whereas mesotrophic assemblages dominated in proximal settings (e.g., Gibbs et al., 2006). The increase in primary productivity observed in coastal areas may have been the result of the injection of nutrients via riverine input. The absence of a nutrient source (e.g., hydrothermal plumes) in open-ocean settings may have caused surface-water starvation. Climatic changes associated with OAE1a and PETM could be responsible for calcareous nannofloral changes, albeit to a lesser extent. The thermal maximum associated with the onset of OAE1a (Menegatti et al., 1998; Jenkyns, 2003) might imply that nannofloral changes derive from warming of deep and intermediate waters, and the consequent to the disruption of the thermocline and elimination of the ecological niche of the deep photic zone nannoconids (Erba, 2004). The warming recorded at high latitudes during the Paleocene-Eocene transition (e.g., Kennett and Stott, 1991) could have partially triggered the shift in abundance of the

Review of calcareous nannofossil changes warm-water taxon Discoaster and the marked decrease of the cold-water taxa Chiasmolithus and Biscutum? sp. (Bralower, 2002; Tremolada and Bralower, 2004). The extinction of Octolithus multiplus, an abundant taxon at northern high latitudes, could have resulted from the abrupt warming recorded at the onset of the PETM. However, the drop of oligotrophic and warm-water taxa such as Nannoconus and Discoaster during the OAE1a and the PETM (only in shelf sites), respectively, may suggest that fertility variations played a more striking role in the composition and abundance of nannofloral assemblages. Alternatively or concurrently, the nannoconid crisis and the decrease in abundance of Discoaster in proximal settings may result from the temporary thermocline disruption caused by extreme temperature. The influence of pCO2 variations on calcareous organisms has been investigated almost exclusively in order to highlight the impact on calcification. Increasing pCO2 levels lowers dissolved [CO32–] (e.g., Feely et al., 2004) and can hinder biomineralization of living marine calcifying organisms such as planktonic foraminifera (e.g., Bijma et al., 1999) and corals (e.g., Gattuso et al., 1998; Kleypas et al., 1999). Likewise, in living coccolithophorids increasing the pCO2 level generally inhibits calcification and may also result in malformation of coccoliths (e.g., Riebesell et al., 2000; Zondervan et al., 2001). In the fossil record, inferred high carbon dioxide concentrations coupled with enhanced fertility correlate with significant reductions in nannofossil carbonate export across the Cretaceous and Jurassic OAEs (Erba and Tremolada, 2004; Erba, 2004; Tremolada et al., 2005). However, it is extremely difficult to link the direct influence of pCO2 levels on changes in abundance and composition of calcareous nannofossil assemblages. Increasing carbon dioxide concentrations affected indirectly calcareous nannofloras and resulted in greenhouse conditions, lower alkalinity, and shallower Calcium Carbonate Compensation Depth (CCD). Greenhouse conditions during OAE1a and the PETM are thought to have been caused by pCO2 levels three to six times (e.g., Erba and Tremolada, 2004) and three to four times (Zachos et al., 2003) higher than pre-event levels, respectively. Humid and warm conditions associated with the “greenhouse world” led to intensified runoff and continental weathering, and increased the stratification of waters, reducing the vertical and latitudinal thermal gradients. Enhanced surface ocean stratification decreases mixing and nutrient input from deeper waters. During OAE1a, nutrient availability was maintained by N2-fixing cyanobacteria (Kuypers et al., 2004; Dumitrescu and Brassell, 2005). The presumed starvation suggested calcareous nannofossil assemblages in open-ocean sites during the PETM could have resulted from the combination of highly stratified waters and limited nutrient input. Diminishing alkalinity and lysocline shoaling lead to the dissolution of previously deposited calcium carbonate (Broecker and Peng, 1982; Walker and Kasting, 1992; Dickens et al., 1997). A significant dissolution event coincident with the onset of the δ13C negative excursion is recorded in both the OAE1a (e.g., Dean, 1981; Premoli Silva et al., 1999; Erba, 2004; Erba and Tremolada, 2004) and the PETM (e.g., Bolle et al., 2000; Raffi et al., 2005;

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Gibbs et al., 2006). However, major nannofloral changes cannot be ascribed to dissolution, but are, conversely, of primary origin. Although several OAE1a sections appeared clearly affected by dissolution and syn- and postdepositional diagenesis (e.g., Erba et al., 1989; Coccioni et al., 1992; Bralower et al., 1994), high abundances of fragile and delicate eutrophic taxa characterize numerous OAE1a sections (e.g., Luciani et al., 2001). In addition, the onset of the “crisis” of the dissolution-resistant taxon Nannoconus predates the deposition of OAE1a black shales (Erba, 1994). Similarly, the decrease in abundance of the robust taxa Discoaster and Fasciculithus, and the increase in dissolution-susceptible mesotrophic taxa recorded in shelf areas during the PETM, cannot be the result of syn- or postdepositional modifications. In general, calcareous nannofloras show small evidence of etching or overgrowth in open-ocean settings at high latitudes (Bralower, 2002; Tremolada and Bralower, 2004). At low latitudes, diagenesis and dissolution significantly affect composition and abundance of calcareous nannofossil communities (e.g., Kahn and Aubry, 2004; Raffi et al., 2005), but results from Shatsky Rise suggest an ecological signal rather than a diagenetic artifact (Gibbs et al., 2006). CONCLUSIONS This study suggests that greenhouse conditions associated with the PETM and the OAE1a had different effects on calcareous nannofloras. The OAE1a nannofossil assemblages suggest eutrophic conditions of surface waters. High primary productivity, triggered by runoff in coastal areas and hydrothermal plumes in open-ocean sites, resulted in low abundances of oligotrophic taxa such as Nannoconus spp. and increasing quantities of meso- to eutrophic species. The PETM generally shows an increase in abundance of oligotrophic taxa such as Discoaster and Fasciculithus and a decrease in mesotrophic taxa during the onset and the peak of the CIE in open-ocean settings. A return to slightly more eutrophic conditions is documented by increasing abundances of mesotrophic species in the upper part of the event at high latitudes. The inferred high-fertility conditions of the PETM interval in proximal settings resulted in calcareous nannofossil assemblages dominated by small-sized mesotrophic taxa. Perhaps, runoff and weathering introduced large quantities of nutrients in proximal settings, whereas a likely negligible or insignificant input of nutrients in open-ocean areas resulted in starvation. Changes in calcareous nannofossil assemblages of both events resulted mainly from variations in fertility of surface waters. Increasing temperature and pCO2 levels had a minor impact on composition and abundance of nannofloral communities. The OAE1a calcareous nannofossils seem to have been influenced by enhanced carbon dioxide concentrations, which hindered biocalcification and decreased the nannofossil paleofluxes. Conversely, the role of pCO2 variations on nannofloral calcification during the PETM is not obvious. However, nannofossil assemblages suggest a greater effect of increased temperature. The increase in abundance of Discoaster at high latitudes may have resulted from

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a drop in primary productivity coupled with climatic conditions that favored the shift in abundance of these nannoliths. Calcareous nannofloras in open-ocean sites at low latitudes could have been indirectly affected by the increase in temperature and pCO2 levels, because greenhouse conditions reduce thermal gradients and enhance stratification of ocean waters. Both events show evidence of lysocline/CCD fluctuations and diagenetic overprint, but overall nannofloral changes suggest a primary signal only slightly altered by diagenesis and dissolution. ACKNOWLEDGMENTS This study was funded by the National Science Foundation (grants OCE-0084032, EAR-9814604, and EAR0120727), Ministero per l’Università e per la Ricerca Scientifica e Tecnologica (Prin-2005044839_001 to I. Premoli Silva), and RPS Energy Ltd. This paper benefited from reviews by W. Wise Jr., S. Jiang, and J. Mutterlose. S. Monechi is warmly thanked for the editorial review. REFERENCES CITED Agnini, C., Muttoni, G., Kent, D.V., and Rio, D., 2006, Eocene biostratigraphy and magnetic stratigraphy from Possagno, Italy: The calcareous nannofossil response to climate variability: Earth and Planetary Science Letters, v. 241, p. 815–830, doi: 10.1016/j.epsl.2005.11.005. Aguado, R., Castro, J.M., Company, M., and de Gea, G.A., 1999, Aptian bioevents—An integrated biostratigraphic analysis of the Almadich Formation, Inner Prebetic Domain, SE Spain: Cretaceous Research, v. 20, p. 663–683, doi: 10.1006/cres.1999.0176. Arthur, M.A., Jenkyns, H.C., Brumsack, H.J., and Schlanger, S.O., 1990, Stratigraphy, geochemistry and palaegeography of organic carbon-rich Cretaceous sequences, in Ginsburg, R.N., and Beaudoin, B., eds., Cretaceous resources, events and rhythms: Dordrecht, Kluwer Academic, p. 75–119. Aubry, M.-P., 1992, Late Paleogene nannoplankton evolution: a tale of climatic deterioration, in Prothero D.R., and Berggren W.A., eds., Eocene– Oligocene climatic and biotic evolution: Princeton, New Jersey, Princeton University Press, p. 272–309. Backman, J., 1986, Late Paleocene to Middle Eocene calcareous nannofossil biochronology from the Shatsky Rise, Walvis Ridge and Italy: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 57, p. 43–59, doi: 10.1016/0031-0182(86)90005-2. Bains, S., Corfield, R.M., and Norris, R.D., 1999, Mechanisms of climate warming at the end of the Paleocene: Science, v. 285, p. 724–727, doi: 10.1126/science.285.5428.724. Bains, S., Norris, R.D., Corfield, R.M., and Faul, K.L., 2000, Termination of global warmth at the Palaeocene/Eocene boundary through productivity feedback: Nature, v. 407, p. 171–174, doi: 10.1038/35025035. Behrenfeld, M.J., and Kolber, Z.S., 1999, Widespread iron limitation of phytoplankton in the South Pacific ocean: Science, v. 283, p. 840–843, doi: 10.1126/science.283.5403.840. Bellanca, A., Erba, E., Neri, R., Premoli Silva, I., Sprovieri, M., Tremolada, F., and Verga, D., 2002, Palaeoceanographic significance of the Tethyan “Livello Selli” (Early Aptian) from the Hybla Formation, northwestern Sicily: Biostratigraphy and high-resolution chemostratigraphic records: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 185, p. 175–196, doi: 10.1016/S0031-0182(02)00299-7. Bijma, J., Spero, J.H., and Lea, D.W., 1999, Reassessing foraminiferal stable isotope geochemistry: Impact of the oceanic carbonate system (experimental results), in Fischer, G., and Wefer, G., eds., Use of proxies in paleoceanography: Examples from the South Atlantic: New York, Springer, p. 498–512. Bolle, M.P., Pardo, A., Hinrichs, K.U., Adatte, T., von Salis, K., Burns, S., Keller, G., and Muzylev, N., 2000, The Paleocene-Eocene transition in the marginal northeastern Tethys (Kazakhstan and Uzbekistan): Inter-

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Printed in the USA

The Geological Society of America Special Paper 424 2007

Ecosystem perturbation caused by a small Late Cretaceous marine impact, Gulf Coastal Plain, USA David T. King Jr. Department of Geology, Auburn University, Auburn, Alabama 36849-5305, USA Lucille W. Petruny Astra-Terra Research, Auburn, Alabama 36831-3323, USA Thornton L. Neathery Neathery and Associates, 1212-H Veterans Parkway, Tuscaloosa, Alabama 35404, USA ABSTRACT The Wetumpka impact event, ca. 83.5 Ma in shallow waters of the northern Gulf of Mexico, caused minor ecosystem perturbations because of the physical effects of a 2.6-gigaton-equivalent impact detonation. The impact event and its consequences had relatively minor, but notable, paleobiologic effects (i.e., preservational effects, biostratigraphic effects, and impact-succession effects). The impact structure served as a local reservoir for an impact-entombed fossil record in two main ways. First, coarse to fine fragments of terrestrial vegetation, probably derived from the adjacent tropical forest, were swept up and incorporated into Wetumpka washback- and surgeback-deposited breccias and sands. Second, intact blocks of target sedimentary units, which contain an internal fossil component of their own, are part of the slump and fallback debris that partially fills the Wetumpka impact structure. Some of these target sedimentary rock blocks include updip sedimentary facies of these formations that no longer exist in outcrops anywhere in the region. In yet another paleoecologic effect, the Wetumpka impact crater apparently functioned as a minor terrestrial (island) ecosystem embedded within the shelfal marine realm for an unknown length of time. As one might expect from the relatively small size of this structure, there is no apparent regional or global biotic extinction event associated with this local catastrophe. Keywords: Wetumpka, crater, impact structure, Alabama, Cretaceous. INTRODUCTION

age of ca. 83.5 Ma (King et al., 2006). The impact-affected area has been mapped using a planetary geology perspective, i.e., employing impact-related, genetic units, as depicted in Figure 2. In previous work, King et al. (2002) reported shocked minerals and iridium enrichment in structurefilling impactites, thus establishing this structure as being of impact origin.

The Wetumpka impact structure is a relatively well preserved marine-target feature of the inner coastal plain of Alabama (Neathery et al., 1976; King et al., 2002, 2003). Wetumpka, which is located at ~32° 31.3′ N, ~86° 10.4′ W (Fig. 1), has a structural diameter of 7.6 km and a stratigraphic

King, D.T., Jr., Petruny, L.W., and Neathery, T.L., 2007, Ecosystem perturbation caused by a small Late Cretaceous marine impact, Gulf Coastal Plain, USA, in Monechi, S., Coccioni, R., and Rampino, M.R., eds., Large Ecosystem Perturbations: Causes and Consequences: Geological Society of America Special Paper 424, p. 97–107, doi: 10.1130/2007.2424(06). For permission to copy, contact [email protected]. ©2007 The Geological Society of America. All rights reserved.

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King et al. crater, which probably washed over the early-formed crater rim and mixed with fallback debris and other impact-related materials (King et al., 2003, 2006; cf. Poag et al., 2004). These washback-surgeback sands and breccias contain distinctive sedimentary structures such as contorted lamination, clay injections, complexly distorted contact geometries, concentrically layered ball-like structures, and crudely graded bedding (cf. Poag et al., 2004; see King et al., 2006). Data supporting conclusions in this paper come from surface investigations at the Wetumpka impact structure (Neathery et al., 1976; Neathery et al., 1997; Nelson, 2000; King et al., 2002) and from subsurface drilling (i.e., two core holes at crater center; discussed in King et al., 2002, 2003). We will address ecosystem perturbation by first discussing physical effects (presenting first a theoretic discussion and noting our petrologic evidence) and subsequently discussing possible paleobiologic effects and our paleontologic evidence for some of these effects. ECOSYSTEM PERTURBATION

Figure 1. Location of the Wetumpka impact structure (dot at end of arrow) in Alabama. The main physiographic/geologic regions of Alabama are indicated.

Impact occurred in shallow shelfal waters of the Late Cretaceous northern Gulf of Mexico at a distance from the adjacent barrier-island shore that has been estimated at ~25 km (Neathery et al., 1997; King et al., 2002; Fig. 3). Paleobathymetry of the impact site is estimated to have been ~30–100 m based on study of optical structures in ostracodes (Puckett, 1991) and the ichnofossil assemblage (Rindsberg, 1990) of the youngest target materials. The impact event targeted a seawater layer, water-saturated wedge of sediment layers, and slightly dipping, underlying crystalline (schist-gneiss) bedrock of the Appalachian Piedmont. The sediment layer consisted of, in reverse stratigraphic order, ~30 m of clayey chalk ooze (Mooreville Chalk), ~30 m of paralic marine sand (Eutaw Formation), and ~60 m of terrestrial clayey sand and gravels (Tuscaloosa Group; Neathery et al., 1997; King et al., 2002). Wetumpka has a two-part crater-filling stratigraphy (Fig. 4) consisting of an upper catastrophic megablock and sand unit, which was deposited by late-stage rim collapse (King et al., 2003, 2006), and a lower impact breccia unit, which is composed of three basic facies: washback- and surgeback-deposited impactite sands and breccias, fallback breccias, and amalgamated slumped target-rock blocks (King et al., 2003, 2006; washback and surgeback terms are sensu Poag et al., 2004). Figure 5 depicts the stages in the structure-filling, stratigraphic sequence. The lower washback-surgeback impactite sands and breccias are interpreted to have formed in part by the collapsing water

Ecosystem perturbation due to an ancient marine impact may be subdivided into physical and paleobiologic effects. Physical effects on the surrounding ecosystem including the adjacent coastline include, in probable order of effect: infrared flashburn; seismic energy; impact blast waves, fallback of ejecta, and tsunami run-up (King et al., 2006). Paleobiologic effects include preservational effects (e.g., rapid burial of dead organisms and reburial of excavated fossils), biostratigraphic effects (e.g., mass death and/or extinction), and impactsuccession effects (e.g., development of new transient ecosystems in the crater depression and on the rim). Physical Effects The Wetumpka impact event is thought to have been caused by the impact of a ~350-m-diameter chondritic asteroid judging from (1) the final crater diameter and (2) the platinum-group trace element assemblage in the structure-filling impactites (King et al., 2002). Using the inferred asteroid diameter and reasonable chondritic density to estimate a probable asteroid mass and assuming that the asteroid velocity was ~20 km/sec (French, 1998), we can use the kinetic energy (KE) formula, KE = ½ mv2,

(1)

i.e., the impact energy from the Wetumpka event (where m is mass and v is velocity), to estimate an impact-derived explosive energy equal to ~1.1 × 1019 Joules or the energy equivalent of ~2.6 × 103 megatons (MT) of TNT (see discussion in French, 1998). The impact energy needed to exceed the “nominal threshold for global disaster” is estimated to be ~3 × 105 MT of TNT (Morrison et al., 1994); therefore, the Wetumpka event clearly falls into the category of a “subglobal” disaster wherein the impact effectively destroyed “the area of a small state” (Morrison et al.,

Figure 2. Generalized geological map of the Wetumpka impact structure (after King et al., 2002). Unit abbreviations: pK—pre-Cretaceous metamorphic rocks; crt—crystalline rim terrain; Ku—Upper Cretaceous target strata; isu—interior sedimentary unit; b—surficial breccia unit; est—extra-structure terrain; m—Mooreville chalk inliers within extrastructure terrain; Qal—Quaternary alluvium. Area of drill sites is marked by asterisk.

Figure 3. Late Cretaceous paleogeography of the Alabama-Georgia state border region (border is dashed). Wetumpka impact (W inside circle) is indicated. After a drawing by W.J. Frazier in Schwimmer (2002).

Figure 4. Impact-structure cross section (west to east) showing interpreted relations between units based upon core drilling and surficial geologic relations, as explained in the text.

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King et al. Infrared Flashburn Emission of electromagnetic waves in the infrared spectrum accompanies impact events as has been shown experimentally, in images of meteorite impacts on the lunar surface and in images taken during the Comet Shoemaker-Levy 9 impact on Jupiter. The fire-ignition potential for such infrared energy is subject to a scaling-law function and can be calculated using a formula written for the purposes of assessing infrared effects of battlefield nuclear weapons (Adushkin and Nemchinov, 1994). It is assumed in this formula that the threshold for ignition is ~100 Joules/cm2 (Adushkin and Nemchinov, 1994). Assuming a clear day (i.e., minimal haze to attenuate the infrared pulse) and an energy budget of 25% of the blast energy (as per King, 1976), we can calculate the infrared burn radius using the AdushkinNemchinov formula for maximum burn radius (Rf), namely Rf = 3E 0.5 km,

Figure 5. Wetumpka drill-core stratigraphy (synthesis of two wells; King et al., 2002, 2003) showing two main units and interpreted origin of each as a stage in crater filling.

1994). If we use the “area of potential forest devastation around a terrestrial impact” as a guide to discerning the impact-affected area from a biotic perspective, the formula given by Morrison et al. (1994) is useful. In that formula, the biotically affected area (Ab) is given approximately by the equation Ab = 104Y 0.666,

(2)

where Ab is the devastated area in hectares and Y is the yield in MT equivalent of TNT. The biotically devastated area for Wetumpka thus calculated (where Y = 2.6 × 103) is ~1.9 × 106 hectares. Assuming essentially circular or “symmetrical” distal paleoecologic effects (cf. Dypvik et al., 2006, who found symmetry to be correct for the large marine crater Mjølnir), the radius of a “disaster circle” would equal ~77 km. This radius is far greater than the distance from the adjacent barrier-island shoreline (cited above), so it can be assumed that a significant segment of the adjacent coastline was involved in the realm of biotic effects from the Wetumpka event. Both shallow seafloor and adjacent coastal lowlands and barrier islands were in the paleoecologic “footprint” of the Wetumpka event.

(3)

where E is the energy budget in MT. The result is an Rf equal to ~76 km. This is about three times the estimated distance to the adjacent shoreline, so the necessary conditions for coastal forest ignition seem to have been in place. Forest combustibles near Wetumpka were likely tropical forest cycads, conifers, angiosperms, and other lush vegetation thriving on Alabama’s low coastal plain, which have been previously described by Mancini (1981). In addition to the direct infrared heating, subsequent atmospheric re-entry of ejected particles that were temporarily sent into low orbits would have secondarily heated the local atmosphere and could have caused additional radiant thermal damage (Toon et al., 1994). It could be anticipated that a substantial quantity of soot and larger burned and partially burned particles would enter the atmosphere as a result of local wildfires. Even though lignite fragments within Wetumpka’s impactites show no flashburn effects such as charcoalization, there is a small component of fine organic material (soot?) in some finegrained intervals of the impactite (Fig. 6; King et al., 2003, 2006). This combination of evidence leads us to speculate that flashburn effects may have been somewhat limited in this instance, and/or that forest fuel supplies may have been limited. Seismic Energy Effects Seismic energy input at Wetumpka may have been ~8.3 × 1018 Joules, if we assume that 75% of the impact kinetic energy was converted to seismic energy as suggested by King (1976). This amount of energy can be related to a surface-wave magnitude equivalent to an earthquake of Richter-scale magnitude ~8.4 to 9 (see Bolt, 1993; King et al., 2006). On a low-lying coastal plain, such as that adjacent to the Wetumpka marine impact, the surfacewave energy of strong seismic waves probably had the potential to bring down high-standing terrestrial vegetation (Bolt, 1993). Physical evidence of seismic effects from Wetumpka may be present in the local stratigraphic section. In a regional synthe-

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Figure 6. Core from well drilled at Wetumpka impact crater center, which shows a lignitic component in impactite sands. Dark material (L) is lignite; dark component in sands (swirled and contorted laminations) is finely divided lignitic material as well. Cored interval: depth is ~118.5 m (upper right) to ~121.8 m (lower left) in the Schroeder well. Scale bar is 15 cm.

sis of Upper Cretaceous stratigraphy, King (1994) pointed out that there appears to be a regionally significant disturbance during Late Cretaceous, the AGCD (“Alabama-Georgia clastic-dike injection”) event, which was also noted by Reinhardt (1980), Frazier (1987), and others. The age of the AGCD event is estimated to have been at least 83 Ma (King, 1994) based upon cross-cutting relationships. The AGCD may be explained by the Wetumpka impact event, which is thought to have occurred ca. 83.5 Ma (King et al., 2006). Impact Blast Waves Impact blast waves have “an abrupt pressure pulse ... followed immediately by a substantial wind” (Toon et al., 1994, p. 794). Whether objects like trees are affected by impact blast waves depends upon peak overpressure, which is defined as the difference between ambient pressure and pressure of the shock front. Peak overpressure equal to 14 kilopascals (kPa) accompanies a maximum wind speed of ~30 m/s. Key peak overpressure of 28 kPa produces maximum wind speed of ~70 m/s, which is equal in velocity to the cyclonic wind speed of a Saffir-Simpson force-5 hurricane (Toon et al., 1994). Thus, in a low-lying coastal forest such as the one near Wetumpka ground zero, near total devastation of the larger biomass (trees) would typically result. Key peak overpressure for a ground impact occurs in an area with maximum radius r, where r = 5.08 (E 0.333) km,

(4)

and E is the adjusted impact energy in MT. E is adjusted according to the equation E = qY,

(5)

where q is an empirically determined constant (0.5 for nuclear weapons) and Y is the kinetic energy yield in MT (Toon et al., 1994). At Wetumpka, a sea-level impact could have set up an atmospheric blast wave that delivered key peak overpressure (28 kPa) at a maximum radius (r) of ~55 km. From the foregoing, we infer that tropical forests on the nearby shoreline would have been profoundly affected by this blast wave, which would have struck the area with straight-line wind speeds somewhat like that of a force-5 hurricane. These winds would have extinguished infrared-ignited fires and may have splintered large amounts of tropical forest wood. This would potentially account for elongate lignite fragments observed in some intervals of the impact breccia within Wetumpka drill cores (Fig. 6; King et al., 2003). Returning air currents, drawn back rapidly by an intensive low pressure at the crater, may have helped transport some splintered wood seaward toward the impact crater and ultimately into the sedimentary structure fill. Fallback of Ejecta The theoretical limit of continuous ejecta from an impact is ~2.35 crater radii from the rim (Melosh, 1989). Therefore, for Wetumpka, the diameter of the continuous ejecta would

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have been ~25 km. Because this area would have been strongly affected by the violent return of displaced seawater, a continuous ejecta blanket is not to be expected in a marine impact such as this (cf. Powars and Bruce, 1999; Ormö and Lindström, 2000; Poag et al., 2004), and none has been found. Discontinuous ejecta potentially fell over a much larger area, including the nearby shoreline and adjacent shallow shelf area. Considering the soft-sediment nature of the upper and more easily excavated target materials, most of the ejected material probably disintegrated into fine particles. Because the crater center was located ~25 km offshore, we would expect only the discontinuous ejecta to fall upon the adjacent land. For this reason, we can say that ballistic sedimentation by such fine ejecta likely did not result in significant comminution of terrestrial vegetation. Discontinuous distal ejecta may exist in the coeval Mooreville Chalk of Dallas County, Alabama (~120 km from Wetumpka), as noted by King and Wylie (1986). They described anomalous intervals in drill cores that consist of sandy marls, ~1 m thick, consisting of a basal scour surface, low-angle cross-lamination, and a graded unit at the top. To date, shocked materials have not been recovered from these intervals so we do not know yet if they are related to Wetumpka. Tsunami Run-Up Terminal wave height (hw) at distance r, in km, from a nuclear explosion in shallow water is given by the general equation hw = 1450 m [(d/r) (Y/1 GT)0.25],

(6)

wherein d is depth of water in m and Y is yield in gigatons (GT; Glasstone and Dolan, 1977; Hills et al., 1994). At distance from shoreline, r, equal to 25 km, and water depth equal to 100 m, according to this equation, the wave height would have been ~7 m. Melosh’s (2003) discussion of recently released naval warfare research by van Dorn et al. (1968) describes experimental nuclear-weapons detonation in shallow marine waters and contains these observations: (1) the amplitude of impact-generated water waves “can never exceed the depth of the ocean” in which they occur; and (2) these waves may be limited to 0.39 times oceanic depth. Thus, we might infer that Wetumpka’s impactor striking in ~30–100 m of water (King et al., 2002) might produce a tsunami-like wave of between ~12 and 39 m at some small distance r from the impact. Amplitudes of tsunami waves decrease with distance (r) from target; however, in this instance, the progressively decreasing water depth toward shore would cause amplitude to increase in a shoreward direction. Low shoaling factors (perhaps <2) associated with impactgenerated wavelengths tend to cause early wave breaking and any Wetumpka tsunami wave breaking in the nearshore realm might be quite limited in its potential tsunami run-up distance (van Dorn et al., 1968; Melosh, 2003).

Considering a relatively flat coastal plain, Hills et al. (1994) estimated a scaled equation for run-up distance, Xmax, for tsunami waves as Xmax = 1000 m [(ho /10 m)1.333],

(7)

where ho is run-up height. Wetumpka’s run-up height, ho, is estimated above to range from ~7 to ~39 m. So, for Wetumpka’s tsunami, which we assume was generated 25 km offshore, tsunami run-up distance, Xmax, would have been between ~0.6 and 6.1 km. On the low end, the estimated distance is hardly enough to cover the nearby barrier island and reach the mainland shore with any significant force or effect. On the high end, barrier islands, coastal lagoons, and probably some mainland forest areas (already burned and affected by earthquakes and the air blast) would be affected. If tsunami effects were involved in transporting woody material from the land to the crater area or to shelf depositional sites, most of that material probably came from the barrier-island realm. PALEOBIOLOGIC EFFECTS Preservational Effects At Wetumpka, there is a small but anomalous lignitic component within some impact breccias and impactite sands (Fig. 6). This lignite is anomalous because very little lignite exists in target rocks (King, 1994; King et al., 2002, 2003). Presence of the lignite, which usually occurs as broken fragments that are angular to subround, is evidence of what we have referred to in previous work as a “fossil reservoir effect” (King et al., 2006). The concept of a terrestrial impact structure, particularly a wettarget or marine impact, as a “reservoir” for preservation of fossils is not difficult to imagine. However, there are not many wellestablished examples (see the discussion of numerous craters in Cockell and Lee, 2002). Terrestrial impact craters that are fossil reservoirs typically have this distinction because of a subsequent lacustrine ecosystem that developed within the crater’s surficial depression (e.g., Ries crater [Pösges and Schieber, 1997]; and Boltysh crater [Gurov et al., 2003]). Wetumpka’s crater fill apparently was not a very good environment for preservation of other coeval marine or terrestrial species—as none have been found— but, as noted above, an allochthonous component of terrestrial organic material (wood fragments and fine debris, now all lignite) was transported by air or water into the crater fill, where it became a component in the impactites. In addition, broken pieces of impacted formations, some of which are several meters across, retain constituent body fossils and trace-fossil assemblages, which occur together in a non-stratigraphic order in both structure-filling units (Fig. 7; King et al., 2002, 2003). Body fossils include various pelecypods (including Exogyra [described by Scott, 1968] in the target Mooreville Chalk). Trace fossil assemblages include (1) Taenidium and related fresh-water traces (described in the target Tuscaloosa Group by Savrda et al., 2000), which occur within

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Figure 7. Core from well drilled at the center of the Wetumpka impact crater, which penetrates a ~6-m Tuscaloosa Group target rock block. Depth of interval is ~164.1 m to 169.7 m (upper left to lower right) in the Schroeder well. Mottled interval in upper part of center three cores contains Taenidium among other bioturbation. Light gray—sand; dark gray—red clays. Scale bar is 15 cm.

intact blocks of Tuscaloosa Group sediments (Fig. 7); and (2) closely related marine trace fossils Planolites and Thalassinoides, which occur within intact blocks of Eutaw Formation and Mooreville Chalk (both marine units described by King, 1994). Except within the blocks of impacted formations mentioned above, there are no trace fossils in Wetumpka’s structurefilling impactites or impact breccias. As noted above, intact blocks of target sedimentary units containing an internal fossil component of their own are part of the slump and fallback debris, i.e., the upper catastrophic megablock and sand unit, which partially fills the Wetumpka impact structure (King et al., 2003, 2006). Two Upper Cretaceous target sedimentary formations (i.e., Eutaw Formation and Mooreville Chalk), which are recognizable as blocks in outcrops near the crater center, no longer crop out in the vicinity of the structure (apparently owing to extensive Tertiary erosion [see Neathery

et al., 1997]). It is worthy of note that these slumped, target sedimentary rock blocks include some updip sedimentary facies of these formations that no longer exist in outcrops anywhere in the region (Neathery et al., 1997; King et al., 2006). Lastly, within the Wetumpka’s surficial structure-filling sediments, stratigraphic mixing or leakage of vertebrate megafossils has been reported. Specifically, a vertebra from a polycotylid pliosaur (Discosaurus), which, judging from some attached chalky matrix, must have originated within the Mooreville Chalk, was recovered from within a target-rock block of sandy Tuscaloosa Group sediment (Thurmond and Jones, 1981). Tuscaloosa is a terrestrial fluvial deposit (Savrda et al., 2000); therefore, in this peculiar instance, a vertebra from a Campanian marine reptile was forced into a block composed of much older Cenomanian fluvial sand deposit owing to cosmic impact (age relations from Neathery et al., 1997).

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Biostratigraphic Effects The inferred chronostratigraphic position of this impact event is within the lowermost ~30 m of the Campanian Mooreville Chalk (Neathery et al., 1997; King et al., 2006). This level is interpreted to be near the local boundary between two planktonic foraminiferal biozones: Dicarinella asymetrica range zone and Globotruncanita elevata interval zone (regional biozones of Mancini et al., 1998). However, we have no evidence that the impact was the cause of any faunal change associated with this biozone boundary. The geochronometric age of this biozone boundary is ca. 83.5 Ma (see global synthesis of biostratigraphy in Haq et al., 1988). On the basis of these relations, Wetumpka’s inferred impact horizon is interpreted to be within the lower Campanian (e.g., see the widely referenced synthesis by Haq et al., 1988). The inferred position of the Wetumpka impact level is also stratigraphically near (1) the ostracode biozone boundary between the Veenia quadrialira and Pterygocythereis cheethami interval zones and (2) the nannoplankton biozone boundary between the Lucianorhabdus cayeuxi and Calculites obscurus interval zones (biozones of Puckett, 1994). A thorough analysis of ostracode and planktonic foraminiferal biostratigraphic distributions in the impact area (Puckett, 1994) showed that there were no coincident extinctions among species of either taxonomic group within the area’s uppermost Santonian or Campanian strata. Regarding megafauna and the Wetumpka impact, we have noted that the inferred impact level is approximately at the same level as the last occurrence horizon of two large oysters, Pyncnodonte aucella and Exogyra upatoiensis (ranges from Sohl and Smith, 1980; King et al., 2006). Whereas this co-disappearance of large species may seem more than coincidental, a similar large oyster in these strata, Exogyra ponderosa, was not so affected. According to Sohl and Smith (1980), Exogyra ponderosa ranges upward to the uppermost Campanian. At present, an exact stratigraphic coincidence between the level of the last occurrence of the two large oysters and the level of Wetumpka impact cannot be demonstrated.

cores (described by King et al., 2002, 2003), which contains a distinctive two-part stratigraphy separated by what we interpret as a paleosol horizon (King et al., 2003, 2006). In the Schroeder drill core, the whole of the interpreted paleosol horizon is 1.65 m thick (Fig. 8) and consists of an upper 28-cm-thick mottled zone of red, clayey siltstone, which grades downward into gray-brown laminated, organic-rich shale. This shale lies directly upon an 8-cm weathered schist clast at the top of the uppermost impact breccia layer in the lower breccia unit. Unfortunately, this interval was not cored in the Reeves well. The 28-cm upper interval is interpreted as a lateritic paleosol, on the basis of our field comparison with lateritic paleosols in nearby terrestrial Upper Cretaceous deposits (cf. Reinhardt and Sigleo, 1983, and Sigleo and Reinhardt, 1988). The iron sesquioxides, mottles, pedotubules, blocky peds, slickensides, and plinthites present in this upper 28-cm interval are among the wellestablished criteria for recognizing paleosols that have developed on sedimentary materials in the U.S. Gulf Coastal Plain (Reinhardt and Sigleo, 1983; Sigleo and Rienhardt, 1988; see also Nettleton et al., 1989, for more on these criteria). Extensive mottling within paleosols, which have developed upon sedimentary materials, is interpreted to represent the effects of plant-root processes during soil development (Mack and James, 1992). The lower gray-brown laminated part of the paleosol interval bears many of the characteristics of an anoxic lacustrine deposit. These characteristics include high organic content (estimated at 15% total organic carbon), fine parallel laminations and lack of

Impact-Succession Effects Calculated crater rim height at Wetumpka, post–transient rim collapse, was likely ~287 m (Neathery et al., 1997; cf. Melosh, 1989), and today the rim averages over 1 km wide at the base (King et al., 2003). This height and width were more than sufficient to exclude local seawater (King et al., 2006). Therefore, in the shallow southern Alabama shelfal sea—at most 100 m deep—the newly formed crater rim would have stood well above sea level. This crystalline rim may very well have been stable for some time after impact and it apparently excluded all or most seawater. Within the rim, there is stratigraphic evidence that at least some structurally uplifted high ground near the crater center was within the subaerial realm. Evidence for a notable time interval of marine-water exclusion is found in one of the drill

Figure 8. Core from well drilled at Wetumpka impact crater center, showing the 1.65-m paleosol interval, which commences at a depth of ~99.6 m in Schroeder well. The upper contact is sharp and iron stained whereas the lower contact is gradual. The lower right core in the righthand box continues in the upper left corner of the left box. Each square on the scale = 1 cm.

Ecosystem perturbation caused by a small Late Cretaceous marine impact bioturbation (visible in the core and in X-ray imagery), and lack of marine micro- or mega- flora and fauna. The absence of salt-crystal impressions in this unit further suggests a fresh-water origin. The gray-brown shale, a very distinctive lithology in the Wetumpka cores, is unlike any of the target formations. In target rock blocks and local undeformed strata, paleosols are found only in Tuscaloosa Group sediments. It is important to note that the Tuscaloosa contains no gray-brown shales or mudstones (King, 1994). The Eutaw Formation contains dark shales, but they are invariably bioturbated and contain marine fossils. The Eutaw has no paleosols in the target area (Frazier, 1987; King, 1994). Thus, we conclude that this interval is not a target rock block. The attitude of the paleosol horizon is also distinctive. Constituent laminations in the gray-brown shale lie, as nearly as can be determined, almost exactly horizontally. Almost without exception, all target rock blocks in both Wetumpka drill cores are oriented in some way so that there is a distinctive dip to laminations and/or other inherent sedimentary structures. Further, internal and/or marginal deformation of target rock blocks is pervasive. In contrast, this interpreted paleosol interval is undeformed except for its upper contact (i.e., the top of the 28-cm lateritic zone), where there is a 1-cm-thick brown layer that is tightly folded as though it were sheared upon emplacement of the overlying megablock. From the evidence above, we interpret the distinctive 1.65-m interval located at the boundary between the lower breccias and the upper catastrophic megablock and sand unit as a paleosol horizon consisting of a former fresh-water lake deposit that was later exposed to subaerial weathering. This would explain why the gray-brown shale grades gradually upward into the 28-cm lateritic zone, which is interpreted as the soil-forming product of a pre–rim collapse interval of tropical weathering. As unlikely as it may seem considering the overall marine setting of this crater, we think that there is sufficient evidence in the core to support the interpretation of a short-lived subaerial ecosystem at or near crater center. The development of Holocene paleosols with similar features in the Gulf Coastal Plain suggests that they can form in 11,000 yr or less (Sigleo and Reinhardt, 1988). In our interpretation, the envisioned crater-center ecosystem persisted for several thousand years until eventual catastrophic rim collapse emplaced the overlying catastrophic megablock and sand unit (Fig. 9; discussed by King et al., 2003, 2006). Wetumpka’s crystalline rim would also have been a site for intensive weathering and probable paleosol development. There is extensive saprolitization (i.e., formation of deep tropical soil C zone units in crystalline bedrock within parts of the rim; Fig. 10), which is indicative of the rim’s significant history of exposure and erosion (cf. Blank, 1978; Reinhardt and Sigleo, 1983; Sigleo and Reinhardt, 1988). Because saprolitization is aided by plant growth, this is further evidence of tropical vegetation development upon the rim during this interval. From the paleosol and saprolite we infer that Wetumpka’s rim and higher parts of the interior probably existed for some time as a terrestrial island ecosystem adjacent to the mainland.

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Figure 9. Artist’s depiction of the catastrophic collapse of the southwestern rim of the Wetumpka crater during Late Cretaceous. View is from the northwest. Based on discussion in King et al. (2002, 2003).

DISCUSSION In a comprehensive review of known and possible biotic effects of impacts, Cockell and Lee (2002) studied several impact structures and described a common succession of paleoecologic effects to be expected from impacts large enough to produce a crater. They listed, in order of occurrence (1) “quasi-complete sterilization” due to impact energy; (2) the phase of hydrothermal biology (from residual heat energy); (3) the phase of “impact succession” (formation and colonization of the crater depression); and (4) the phase of ecological assimilation (upon filling or erosion of the structure). This sequence of effects is not so clearly defined at Wetumpka, perhaps because of the marine setting and the modest size of the impact event. As noted above, the phase of “quasi-complete sterilization” affected the Gulf Coast ecosystem over “the area of a small state” (Morrison et al., 1994), with the radius of the “disaster circle” equal to ~77 km. This “subglobal” disaster is not reflected in the paleontological record, as best we can determine. Until there is a better age constraint upon the Wetumpka impact itself and the ages of the biozone boundaries noted above, any correlation remains speculative. Neither is the phase of hydrothermal biology clearly displayed at Wetumpka, as it is at other craters (Cockell and Lee, 2002), perhaps because of the relatively small size of this impact and it shallow marine setting. The phase of “impact succession” is shown at Wetumpka in an indirect yet clear way. Instead of a fossil record of succession, paleosols and saprolites—the development of which is greatly aided by vegetative cover—are present within and on the rim of the crater, respectively. These sites were formerly small post-impact terrestrial ecosystems within the larger context of the formerly marine realm. The phase of ecological assimilation at Wetumpka, Cockell and Lee’s (2002) final phase, continues today at Wetumpka.

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Figure 10. Outcrops on northwestern part of Wetumpka’s rim showing Late Cretaceous weathering features. (A) Upper Cretaceous saprolite overlies hard crystalline rim; height of outcrop is ~10 m. (B) Deeply weathered zone of crystalline rim is overlain by Upper Cretaceous saprolite, which is truncated by a coarse, white littoral sand deposit; height of outcrop is ~15 m.

The structure has been filled (or partially filled) with sediment since Late Cretaceous and is in the process of being exhumed, in postglacial times, from a Quaternary alluvial sedimentary fill (Neathery et al., 1976; Reinhardt et al., 1979; Neathery et al., 1997; Nelson, 2000). During Holocene, lithic inhomogeneity related to the geology of Wetumpka’s rim and crater floor has had a strong effect upon local drainage (i.e., annular drainage of local streams and the Holocene course of the Coosa River in the crater’s vicinity; Neathery et al., 1976; Nelson, 2000). ACKNOWLEDGMENTS The authors thank the City of Wetumpka and the City-County Crater Commission for its continued support of our research. Funding was provided by contributors to the Wetumpka Impact Crater Fund at Auburn University. Drilling at the Wetumpka impact crater was provided by Vulcan Materials Company, Birmingham, Alabama. We thank local landowners for access to their property for exploration and drilling. Figure 9 was provided by the City of Wetumpka from among commissioned artworks by Jerry Armstrong, Atlanta, Georgia. This paper was improved by reviewers David Powars and Maurits Lindström. REFERENCES CITED Adushkin, V.V., and Nemchinov, I.V., 1994, Consequences of impacts of cosmic bodies on the surface of the Earth, in Geherls, T., ed., Hazards due to comets and asteroids: Tucson, University of Arizona Press, p. 721–778. Blank, H.R., 1978, Fossil laterite on bedrock in Brooklyn, New York: Geology, v. 6, p. 21–24, doi: 10.1130/0091-7613(1978)6<21:FLOBIB>2.0.CO;2. Bolt, B.A., 1993, Earthquakes: New York, W.H. Freeman, 331 p. Cockell, C.S., and Lee, P., 2002, The biology of impact craters—A review: Biological Reviews, v. 77, p. 219–310.

Dypvik, H., Smelnor, M., Sandbakken, P.T., Salvigsen, O., and Kalleson, E., 2006, Traces of the marine Mjølnir impact event: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 241, p. 621–636. Frazier, W.J., 1987, A guide to facies stratigraphy of the Eutaw Formation in western Georgia and eastern Alabama, in Frazier, W.J., and Hanley, T.B., eds., Geology of the Fall Line: A field guide to structure and petrology of the Uchee Belt and facies stratigraphy of the Eutaw Formation in southwestern Georgia and adjacent Alabama: Atlanta, Georgia Geological Society, p. B1–B25. French, B.M., 1998, Traces of catastrophe: A handbook of shock-metamorphic effects in terrestrial meteorite impact structures: Houston, Texas, Lunar and Planetary Institute Contribution No. 954, 120 p. Glasstone, S., and Dolan, P.J., 1977, The effects of nuclear weapons (third edition): Washington, D.C., U.S. Government Printing Office. Gurov, E.P., Kelley, S.P., and Koeberl, C., 2003, Ejecta of the Boltysh impact crater in the Ukranian shield, in Koeberl, C., and Martínez-Ruiz, F., eds., Impact markers in the stratigraphic record: Berlin, Springer, p. 179–202. Haq, B.U., Hardenbol, J., and Vail, P.R., 1988, Mesozoic and Cenozoic chronostratigraphy and eustatic cycles, in Wilgus, C.K., Hastings, B.S., Kendall, G.C.St.C., Posamentier, H.W., Ross, C.A., and van Wagoner, J.C., eds., Sea-level changes: An integrated approach: Society of Economic Paleontologists and Mineralogists Special Publication 42, p. 71–108. Hills, J.G., Nemchinov, I.V., Popov, S.P., and Teterev, A., 1994, Tsunami generated by small asteroid impacts, in Geherls, T., ed., Hazards due to comets and asteroids: Tucson, University of Arizona Press, p. 779–789. King, D.T., Jr., 1994, Upper Cretaceous depositional sequences in the Alabama Gulf Coastal Plain: Their characteristics, origin, and constituent clastic aquifers: Journal of Sedimentary Research, v. B64, p. 258–265. King, D.T., Jr., and Wylie, J.A., 1986, Sedimentary facies and sea-level cycles of the Upper Cretaceous Mooreville Chalk, central Alabama: Gulf Coast Association of Geological Societies Transactions, v. 36, p. 489–495. King, D.T., Jr., Neathery, T.L., Petruny, L.W., Koeberl, C., and Hames, W.E., 2002, Shallow marine-impact origin for the Wetumpka structure (Alabama, USA): Earth and Planetary Science Letters, v. 202, p. 541–549, doi: 10.1016/S0012-821X(02)00803-8. King, D.T., Jr., Neathery, T.L., and Petruny, L.W., 2003, Crater-filling sediments of the Wetumpka marine-target impact crater (Alabama, USA), in Dypvik. H., Burchell, M.J., and Claeys, P., eds., Cratering in marine environments and on ice: Berlin, Springer, p. 97–113. King, D.T., Jr., Petruny, L.W., and Neathery, T.L., 2006, Paleobiotic effects of the Late Cretaceous Wetumpka marine impact, a 7.6 km diameter impact

Ecosystem perturbation caused by a small Late Cretaceous marine impact structure, Gulf Coastal Plain, USA, in Cockell, C.S., ed., Biotic effects of impacts: Berlin, Springer, p. 121–142. King, E.A., Jr., 1976, Space geology, an introduction: New York, Wiley and Sons, 349 p. Mack, G.H., and James, W.C., 1992, Paleosols for sedimentologists: Boulder, Colorado, Geological Society of America, Short Course Notes, 127 p. Mancini, E.A., 1981, Paleobotanical studies in Alabama: 1845–1980: Tuscaloosa, Geological Survey of Alabama Circular 106, 21 p. Mancini, E.A., Puckett, T.M., Parcell, W.C., Crow, C.J., and Smith, C.C., 1998, Sequence stratigraphy and biostratigraphy of Upper Cretaceous strata of the Alabama Coastal Plain, in Mancini, E.A., and Puckett, T.M., eds., Sequence stratigraphy and biostratigraphy of Upper Cretaceous strata of the Alabama Coastal Plain: Tuscaloosa, Alabama Geological Society Guidebook 35, p.1–12. Melosh, H.J., 1989, Impact cratering, a geologic process: New York, Oxford University Press, 245 p. Melosh, H.J., 2003, Impact-generated tsunamis: An over-rated hazard: Lunar and Planetary Science Conference, 34th annual meeting, abstract 2013. Morrison, D., Chapman, C.R., and Slovic, P., 1994, The impact hazard, in Geherls, T., ed., Hazards due to comets and asteroids: Tucson, University of Arizona Press, p. 59–92. Neathery, T.L., Bentley, R.D., and Lines, G.C., 1976, Cryptoexplosive structure near Wetumpka, Alabama: Geological Society of America Bulletin, v. 87, p. 567–573, doi: 10.1130/0016-7606(1976)87<567:CSNWA>2.0.CO;2. Neathery, T.L., King, D.T., Jr., and Wolf, L.W., eds., 1997, The Wetumpka impact structure and related features: Tuscaloosa, Alabama Geological Society Guidebook 34c, p. 25–56. Nelson, A.I., 2000, Geological mapping of the Wetumpka impact crater area, Elmore County, Alabama [M.S. thesis]: Auburn, Alabama, Auburn University, 187 p. Nettleton, W.D., Gamble, E.E., Allen, B.L., Borst, G., and Peterson, F.F., 1989, Relict soils of subtropical regions of the United States, in Bronger, A., and Catt, J.A., eds., Paleopedology, nature and application of paleosols: Catena, suppl. 16, p. 59–94. Ormö, J., and Lindström, M., 2000, When a cosmic impact strikes the sea bed: Geological Magazine, v. 137, p. 67–80, doi: 10.1017/ S0016756800003538. Poag, C.W., Koeberl, C., and Reimold, U.W., 2004, The Chesapeake Bay crater: Geology and geophysics of a late Eocene submarine impact structure: Berlin, Springer, 522 p. Pösges, G., and Schieber, M., eds., 1997, The Ries Crater Museum Nördlingen: Munich, Bavarian Academy for Teacher Training, Academy Bulletin 253, 80 p. Powars, D.S., and Bruce, T.S., 1999, The effects of the Chesapeake Bay impact crater on the geological framework and correlation of hydrogeologic units of the lower York-James peninsula, Virginia: U.S. Geological Survey Professional Paper 1612, 82 p.

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Puckett, T.M., 1991, Absolute paleobathymetry of Upper Cretaceous chalks based on ostracodes: Evidence from the Demopolis chalk (Campanian– Maastrichtian) of the northern Gulf Coastal Plain: Geology, v. 19, p. 449– 452, doi: 10.1130/0091-7613(1991)019<0449:APOUCC>2.3.CO;2. Puckett, T.M., 1994, Planktonic foraminiferal and ostracode biostratigraphy of upper Santonian through lower Maastrichtian strata in central Alabama: Gulf Coast Association of Geological Societies Transactions, v. 44, p. 585–595. Reinhardt, J., 1980, Upper Cretaceous stratigraphy and depositional environments, in Frey, R.W., ed., Excursions in southeastern geology, Volume 2: Alexandria, Virginia, American Geological Institute, p. 386–392. Reinhardt, J., and Sigleo, W.R., 1983, Mesozoic paleosols: Examples from the Chattahoochee River valley, in Carrington, T.J., ed., Current studies of Cretaceous formations in eastern Alabama and western Georgia: Tuscaloosa, Alabama Geological Society Guidebook 20, p. 3–10. Reinhardt, J., Prowell, D.C., Christopher, R.A., and Markewich, H.W., 1979, Cenozoic tectonics in the southeast: Evidence from sediments near Warm Springs, Georgia: Geological Society of America Abstracts with Programs, v. 11, no. 4, p. 209. Rindsberg, A.W., 1990, Cretaceous trace fossils in Alabama chalks: Tuscaloosa, Alabama Geological Society Guidebook 26, p. 111–119. Savrda, C.E., Blanton-Hooks, A.D., Collier, J.W., Drake, R.A., Graves, R.L., Hall, A.G., Nelson, A.I., Slone, J.C., Williams, D.D., and Wood, H.A.R., 2000, Taenidium and associated ichnofossils in fluvial deposits, Tuscaloosa Formation, eastern Alabama, southeastern USA: Ichnos, v. 7, p. 227–242. Schwimmer, D.R., 2002, King of the crocodylians: The paleobiology of Deinosuchus: Bloomington, Indiana University Press, 220 p. Scott, J.C., et al., 1968, Facies changes in the Selma Group in central and eastern Alabama: Tuscaloosa, Alabama Geological Society Guidebook 14, 69 p. Sigleo, W.R., and Reinhardt, J., 1988, Paleosols from some Cretaceous environments in the southeastern United States, in Reinhardt, J., and Sigleo, W.R., eds., Paleosols and weathering through geologic time: Principles and applications: Geological Society of America Special Paper 216, p. 123–142. Sohl, N.F., and Smith, C.C., 1980, Notes on Cretaceous biostratigraphy, in Frey, R.W., ed., Excursions in southeastern geology, Volume 2: Alexandria, Virginia, American Geological Institute, p. 392–402. Thurmond, J.T., and Jones, D.E., 1981, Fossil vertebrates of Alabama: Tuscaloosa, University of Alabama Press, 242 p. Toon, O.B., Zahnle, K., Turco, R.P., and Covey, C., 1994, Environmental perturbations caused by impacts, in Geherls, T., ed., Hazards due to comets and asteroids: Tucson, University of Arizona Press, p. 791–827. van Dorn, W.G., LeMéhauté, B., and Hwang, L.-S., 1968, Handbook of explosion-generated water waves, Volume 1, State of the art: Pasadena, California, Tetra Tech. MANUSCRIPT ACCEPTED BY THE SOCIETY 20 NOVEMBER 2006

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The Geological Society of America Special Paper 424 2007

Chemostratigraphy of Frasnian-Famennian transition: Possibility of methane hydrate dissociation leading to mass extinction Mohammad Hossein Mahmudy Gharaie* Department of Geology, Ferdowsi University of Mashhad, 91775-1436, Mashhad, Iran Ryo Matsumoto Department of Earth and Planetary Science, University of Tokyo, Hongo 7-3-1, 113-0033, Tokyo, Japan Grzegorz Racki Department of Ecosystem Stratigraphy, University of Silesia, Będzińska str. 60, 41-200 Sosnowiec, Poland Yoshitaka Kakuwa Graduate School of Arts and Sciences, University of Tokyo, Komaba 3-8-1, 153-8902, Tokyo, Japan

ABSTRACT Various scenarios have been proposed to explain the Late Devonian mass extinction, foremost among which are bolide impact and sea-level fall. We hereby propose a gas hydrate-induced model based on detailed geochemical and sedimentological data. The period of enhanced organic carbon burial in Iran, in south China, and in subpolar Urals corresponds to a brief negative δ13C excursion of 3.5‰ at the Frasnian-Famennian (F-F) transition. Prior to this event, oceanic δ13C increased for a period of several million years. However, major perturbations of the carbon geochemical cycle, and corresponding sharp and strong negative spikes of δ13C, which require a large input of isotopic light carbon into the ocean, also characterize the boundary horizons. Oxygen isotope ratios show negative excursions of 1.7‰ in south China and 4.1‰ in subpolar Urals that parallel the negative excursions in δ13C values. Synchronous negative spikes of δ18O are likely to imply a rapid increase of ocean temperature. We propose that the F-F boundary event was ultimately caused by voluminous and abrupt release of methane from marine gas hydrate into the ocean and atmosphere to trigger rapid global warming. Assuming that the total amount of inorganic carbon of the Devonian ocean was 40,000 gigatons (Gt) and δ13C of gas hydrate methane was −80‰, only 2600 Gt carbon from the total amount of 10,000 Gt gas hydrate carbon could have changed the oceanic δ13C values from +1‰ to −3‰, the observed magnitude of the F-F boundary excursion. Therefore only ~26% of the gas hydrate could have triggered the boundary events. Widespread rift-related, basaltic volcanism along eastern Laurussia and northern Gond*[email protected]

Gharaie, M.H.M., Matsumoto, R., Racki, G., and Kakuwa, Y., 2007, Chemostratigraphy of Frasnian-Famennian transition: Possibility of methane hydrate dissociation leading to mass extinction, in Monechi, S., Coccioni, R., and Rampino, M.R., eds., Large Ecosystem Perturbations: Causes and Consequences: Geological Society of America Special Paper 424, p. 109–125, doi: 10.1130/2007.2424(07). For permission to copy, contact [email protected]. ©2007 The Geological Society of America. All rights reserved.

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Gharaie et al. wana during the middle Late Devonian is believed to have contributed greatly to the global warming surrounding the F-F boundary, which in turn would have triggered massive dissociation of methane hydrate, especially if paired with intensive igneous and tectonic activity and rapid sea-level fall. Keywords: Frasnian-Famennian boundary, mass extinction, isotope geochemistry, methane hydrate.

INTRODUCTION The Frasnian-Famennian (F-F) interval of the Late Devonian has attracted much attention because of its link to one of the five major mass extinctions in the Phanerozoic (Raup and Sepkoski, 1982). Worldwide boundary strata are often characterized by black laminated shale, suggesting low oxygenation and an increased rate of organic carbon burial. This anoxia crisis is known as the Kellwasser (KW) Event (House, 1985). Some studies have focused on paleontological arguments (extinction and reduced origination rates) whereas others have focused on geochemical data such as iridium content, isotopic compositions of C, O, and S (Buggish, 1991), and other trace elements and rare earth elements (Claeys et al., 1996). The cumulative effect of a long series of either terrestrial or extraterrestrial events has been suggested as the cause of the Kellwasser Event (Sandberg et al., 1988, 1992). Impact-triggered environmental changes, such as global cooling, oceanic overturn, and hydrothermal activity, were often viewed as a direct killing factor (e.g., Geldsetzer et al., 1993; Bai et al., 1994; McGhee, 1996). On the other hand, multicausal, exclusively terrestrial mechanisms have received a strong consideration, as reviewed by Racki (2005). For example, repeated, possibly autocyclic co-occurrences of sealevel oscillations, anoxic conditions, and climatic changes have been regarded as effective causes (e.g., Schindler, 1990; Buggisch,

1991; Becker and House, 1994; Joachimski and Buggisch, 1996, 2002; Hallam and Wignall, 1997; Racki, 1998; Chen et al., 2002, 2005; Joachimski et al., 2001, 2004; Averbuch et al., 2005). Although our understanding of F-F boundary event(s) has recently improved through study of marine carbonate successions from North America, Australia (Canning Basin), and Europe (Poland, Belgium, France; Racki and House, ed., 2002), comparatively little has been documented in Africa and Asia and so the course of environmental change there is poorly known. In addition, the mechanism of the F-F mass extinction still remains largely conjectural. The purpose of this paper is to present the results of an integrated geochemical and sedimentological study of widely exposed Upper Devonian successions in three regions of the Tethyan realm (Fig. 1), in order to provide new insights into environmental perturbation leading to the F-F boundary bio-event. Original data from Iran and south China are compared with reinterpreted geochemical results from subpolar Urals (Yudina et al., 2002). GEOLOGICAL SETTING The Upper Devonian sections of the F-F boundary strata have been studied in central Iran, south China, and the polar Urals. Iran, south China, and subpolar Urals were on the margins of the semi-restricted Paleotethys Ocean during the Devonian.

Paleo Tethys

Figure 1. Late Devonian paleogeography reconstructed from Scotese and McKerrow (1990). Iran plate (1), South China plate (2), and Pechora plate (3) face onto semi-restricted Paleotethys Ocean.

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et al., 1999; Dastanpour and Aftabi, 2002) as Upper Devonian. The Bahram Formation consists of variable lithologies and is characterized by cyclic sedimentation of bedded, dark, fossiliferous limestone, gray shale, and quartzose sandstone. Limestone dominates the middle part of the Bahram Formation and is late Frasnian to early Famennian in age. The F-F boundary occurs in a dark gray-black shale unit with thin intercalations of dark gray micritic limestones (Fig. 3).

Iran

South China

Hutk Section Devonian deposits are widely distributed in Iran (Huckriede et al., 1962; Gaetani, 1965; Stocklin et al., 1965; Davodzadeh and Weber-Diffenbach, 1987; Alavi Naini, 1993; Ghavidel Syooki, 1994; Dastanpour, 1996; Ashouri, 1997; Wendt et al., 1997, 2002; Yazdi, 1999, 2001), but the exact stratigraphic position of the F-F boundary has not been clearly defined in most successions because of the scarcity of deepwater conodonts. The Hutk section (Huckriede et al., 1962) is one of the best-dated sections in southeast Iran (Golshani et al., 1973; Janvier, 1974; Brice et al., 1999; Wendt et al., 1997, 2002). Fossiliferous carbonate units of the Bahram Formation in southeastern Iran (Stocklin et al., 1965) on the recumbent anticline, east of Hutk village (Fig. 2), Kerman Province, are well dated on the basis of conodont biozones (Wendt et al., 1997, 2002) and brachiopods (Rashidi, 1993; Dastanpour, 1996; Brice

Liujing Section The Liujing section in Hengxian, Guangxi, is a well-known Devonian section in south China (Fig. 4). It is near the Liujing railway station 60 km east of Nanning. The section is made up of the Gubi Formation and the Rongxian Formation. The Gubi Formation, late Givetian and Frasnian in age, is composed of platform margin deposits. The uppermost part is characterized by dark gray lenticular micritic limestone intercalated with very thin bedded or laminated calcareous claystones. The base of the Rongxian Formation is marked by ~20 cm of thick, light gray micritic limestone. This basal limestone is overlain by very thickly bedded, lightcolored brecciated limestone with few fossils of Famennian age (Fig. 4B). Brachiopods, corals, and algae have been found in the upper part of the Rongxian Formation (Kuang et al., 1989). The F-F boundary transition, a 15-cm interval sandwiched between bedded and brecciated limestones, was defined by the occur-

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Both the Iranian Plate and the South China Block were located in the Southern Hemisphere. The Iranian Plate was situated on the northern margin of Gondwana close to 30° S (Scotese and McKerrow, 1990; Stampfli et al., 2002), and the South China Block along the southeast margin of the Paleotethys at ~20–10° S. The subpolar Urals, on the other hand, were within the Pechora Plate, in the Northern Hemisphere at a latitude close to 15° N on the eastern margin of the Laurussia (Fig. 1).

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Figure 2. Hutk section locality in Iran. Geological map of Hutk area, north of Kerman, with the location of studied section (shown by star).

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Figure 3. Stratigraphy, carbon and oxygen isotope variations, and strontium isotope ratios in the Hutk section. The Frasnian-Famennian boundary is based on conodont zonation after Gharaie (2002) and Wendt et al. (2004).

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Figure 4. (A) Locality of the Liujing section in Guangxi province, south China. (B) Liujing stratigraphic columnar section. Photograph shows position of the F-F boundary. Length of marker (lower arrow) is 14 cm.

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rence of the well-preserved conodonts Palmatolepis linguiformis (latest Frasnian) and Pa. triangularis (earliest Famennian). Micritic limestone and laminated calcareous claystone from the upper part of Gubi formation and the lowermost Rongxian Formation were collected for geochemical analysis. Subpolar Urals Syv`yu River Section The Upper Devonian deposits in the West Urals structural zone are exposed along the right bank of the Syv`yu River in several outcrops. The Syv`yu River section is located ~38 km upstream from its junction with the Kozhym River (Fig. 5C), and comprises undisturbed well-bedded deposits of Vorota Formation (Yudina et al., 2002). The Vorota Formation represents an intrashelf basin fill with mostly bituminous sediments such as siliceous, terrigenous, and carbonate sets of the Frasnian to lower Famennian (Fig. 6). Conodont biostratigraphy was established by Yudina (1989) and Savage and Yudina (1999). The F-F transition is mostly argillaceous (see Yudina et al., 2002, for detailed introduction). GEOCHEMICAL RESULTS Stable isotopic composition of carbon, oxygen and strontium, and abundance of trace elements such as Cu, Mo, and Zn are considered geochemical proxies of environmental perturbation in the Paleotethyan realm. Carbon Isotope The δ13Ccarb of the Hutk section exhibits a strong negative carbon isotope event (CIE) within the boundary strata (Fig. 3). δ13C in the lower part of the Hutk section (Frasnian strata)

ranges from 0.41‰ to 1.5‰ (PDB), with average of ~1‰ (a typical Upper Devonian background value, after Joachimski and Buggisch, 1996; see Figure 8 in Yudina et al., 2002). The δ13C value of carbonates within the F-F boundary interval ranges from +1‰ to −3‰ and exhibits negative excursions. The negative shift is 4‰ in magnitude. In the upper Famennian strata, δ13Ccarb averages ~1‰, oscillating within a narrow range; these values are different from those at the F-F boundary transition, and the lower strata. The δ13C values show a negative excursion of 1.3‰ at the F-F boundary (linguiformis to triangularis conodont biozones) of the uppermost Gubi Formation and the lower Rongxian Formation of the Liujing section in south China (Fig. 7). The sharp negative excursion at Liujing, notably recorded in limestone succession lacking anoxic facies, is likely to be correlated with the δ13C negative excursion of the Hutk section. The negative δ13C excursion is confirmed by data documented in other localities in south China (Chen et al., 2002), and the high-resolution C-isotopic curves reveal negative excursions in the latest Frasnian in both carbonates and organic carbon (Chen et al., 2005). The carbon isotope record of the Syv`yu River section shows a significant increase in δ13C below the F-F boundary from 0 to 3.5‰, in the supposed Upper KW level. High δ13C values continue throughout the lower Pa. triangularis zone but values gradually approach the background level of 1‰ in the top of the sequence (Fig. 6). The isotopic values measured in the intercalated micritic and shaly deposits show some scatter, but significantly two consecutive samples in the thin-bedded, lower part of the Upper KW level show a strong negative δ13Ccarb anomaly of 2.5‰. A similar negative bias is detected in six pilot samples from an alleged Lower KW level. Overall the carbon isotope record in the Syv`yu River section shows the increasing δ13C pattern over a longer period for the Upper KW level

Figure 5. Syv`yu River section locality in subpolar Urals (after Yudina et al., 2002). The location of the Syv`yu River section is shown by star in sketch map C.

SAMPLES

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Figure 6. Stable isotope geochemistry for the upper Frasnian and lower Famennian in the Syv`yu River section (after Yudina et al., 2002); three “omitted samples” in Figure 8 of Yudina et al. are distinguished by very light carbon and oxygen isotopic enrichments and thought by Yudina et al. (2002) to be seriously biased by diagenesis. Thus, the interval on the simplified curve (i.e., without the three points) is marked as a broken line to emphasize the lacking record. G, Y, and B in sedimentary features column refer to colors: G—gray; Y—yellow; B—black. See Figure 8 in Yudina et al. (2002) for more detail.

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Figure 7. Stratigraphy, stable isotope geochemistry, and redox-sensitive elements for the upper Frasnian and lower Famennian in the Liujing section.

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Chemostratigraphy of Frasnian-Famennian transition

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(Joachimski et al., 2002); however, it confirms the abrupt negative excursion around the F-F boundary, which is exhibited in the Iran and south China sections. Also, in Poland (Kowala section, Holy Cross Mountains; see Figures 8 and 9 in Joachimski et al., 2001) a 1–2‰ decline of δ13C values is recorded before the major positive excursion across the F-F boundary both in inorganic and in some photoautotroph biomarkers.

V have been concentrated in a 10-cm-thick micritic limestone of the Liujing section that contains brown cubic molds of pyrite pseudomorphs. The (Cu+Mo)/Zn index can be used as an indicator for redox of seawater. It exhibits a high ratio at the boundary time span in the Liujing section (Fig. 7).

Oxygen Isotope

Cause of the Negative Excursion of δ13C

The large scatter of δ18O values ranging from –10‰ to +5‰—and with an irregular pattern in the Hutk section (Fig. 3)— may reflect postdepositional processes. The Iran plate is located in Alpine-Himalayan orogenic belt; therefore, thermal effects on oxygen isotopes may have occurred during tectonic activities here, as well as in the polar Urals, which are situated at the active continental margin (see Figure 4–5 in Khain and Seslavinsky, 1996). The Liujing section, on the other hand, is located on a passive plate and is therefore less affected by orogenic movement. The most reliable δ18O record of the Liujing section shows an abrupt drop from –4.2‰ down to –5.9‰ at the CIE horizon (Fig. 7), suggesting that the CIE was accompanied by an increase of ocean temperatures. The oxygen isotope variation in the Syv`yu River section exhibits a lighter record than the Liujing section on average, remaining stable at around −6‰ (Fig. 6), and this may confirm either a primary origin for the isotopic compositions or pervasive diagenetic overprint. However, the significant short-lived negative δ18O excursion down to −10.4‰ is similarly restricted in the Syv`yu River section to the 20-cm-thick interval just at the F-F boundary.

Marine anoxia is commonly accepted as the causal event for the middle Late Devonian mass extinction, because two discrete transgressive-anoxic episodes (Lower and Upper Kellwasser Events) correspond to the extinction horizons (e.g., House, 1985; Sandberg et al., 1988, 2002; Schindler, 1990; Buggisch, 1991; Walliser, 1996). As discussed by Joachimski and Buggisch (1993, 1996, 2002), Joachimski et al. (2001, 2002), and Chen et al. (2005), the late Frasnian and early Famennian inorganic and organic carbon isotope records usually show two positive worldwide δ13C excursions during the anoxic events, with an amplitude of ~3‰. Consequently, a relatively large negative scatter in δ13Ccarb values at the F-F transition, exemplified in micrite-clayey matrix in the Syv`yu River section (Yudina et al., 2002) and the Kowala section in Poland (Joachimski et al., 2001), was invariably explained, following Coniglio’s (1989) work, by shallowburial lithification and recrystallization of the carbonate muds during anaerobic oxidation of organic matter and input of δ13Cdepleted bicarbonate (Joachimski et al., 2001, 2002). However, some authors (e.g., Wang et al., 1991, 1996; Bratton, 1999; Chen et al., 2002; Matsumoto et al., 2002) suggested a primary origin of the negative δ13Ccarb signals—a view buttressed by δ13Corg data (Chen et al., 2005)—even if the inexact correlation between biostratigraphy and chemostratigraphy needs additional explanation (perhaps this is the result of dating errors or diachronous C-isotopic variation during sea-level change; see Panchuk et al., 2006). This hypothesis is explored herein in the context of catastrophic methane release suggested previously by Bratton (1999), Chen et al. (2002), and Matsumoto et al. (2002). Cathodoluminescent microscopy reveals less diagenesis (recrystallization) in the micritic limestone of the Liujing section, which is significant and may reflect the original isotopic composition of the seawater. Moreover, the occurrence of negative δ13Ccarb spikes in three different localities from three separate continents (Fig. 1), simultaneous with other geochemical signatures at the F-F boundary such as δ18O and concentration of redox-sensitive elements, are more likely to have been caused by the same phenomena. In the other words, diagenesis/recrystallization is probably not considered important at the F-F boundary. Large, light carbon inputs that might have caused the major δ13C negative excursion have been considered (see review in Berner, 2002), including: 1. A drop in the global Corg burial rate due to a sudden drop in primary productivity; 2. Rapid oxidation of living biomass;

Strontium Isotopes Ratio The estimated path of seawater 87Sr/86Sr during the F-F interval is based on analyses of micritic limestone from the Hutk and Liujing sections. The 87Sr/86Sr ratio shows less variation during the lower and middle Frasnian, and it rises from 0.70845 to 0.70869 in the F-F boundary interval in the Hutk section (Fig. 3). The 87Sr/86Sr remains high and shows minor variations from 0.70865 to 0.70870 through the late Frasnian, and then it finally falls to 0.70835 through the F-F transition and keeps low values through the middle to late Famennian. The 87Sr/86Sr data of the Liujing section (Fig. 7) show a pattern similar to that of the Hutk section, but more refined data from this region show a noticeable negative-to-positive Sr-isotopic pattern in both of the broad KW intervals (Chen et al., 2005). Redox-Sensitive Elements Redox-sensitive metals such as U, V, Mo, As, Ni, and Cu are usually used as indicators of redox conditions. Micritic limestones from the Liujing section contain high concentrations of redox-sensitive elements in certain horizons, which coincide with the negative δ13C excursion at the F-F boundary. Mo, As, U, and

DISCUSSION

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3. Oxidation of a considerable volume (several thousand Gt) of “buried organic materials” related to tectonic disturbance and climatic change; 4. Influx into the basin of isotopically light volcanogenic CO2 from volcanic activity; 5. Massive dissociation of methane hydrates buried at shallow depth beneath cold and/or deep seas as repositories of isotopically light carbon (Dickens et al., 1997); and 6. Sudden upwelling or turnover of stratified seawater. According to the Spitzy and Degens (1985) formula, and assuming an oceanic dissolved inorganic carbon reservoir as large as the present-day one, “drop in organic activity” and “rapid oxidation of living biomass” are not possible as causes of the δ13C negative excursion at the F-F boundary. The degree of “oxidation of buried organic carbon” cannot reliably be estimated. The sharpness of the negative excursion indicates that the event of light carbon release was abrupt and of short duration, probably about tens of thousands years (roughly estimated from the F-F boundary; also see Chen and Tucker, 2003). But the oxidation of buried organic carbon by exposure to air or water seems to have taken much longer. This kind of mechanism could be effective in explaining a δ13C excursion event of much longer duration. Calculation based on Spitzy and Degens (1985) formula and the same assumptions indicates that 50,000 Gt = 50,000* 1015g of carbon of volcanic and/or hydrothermal origin (δ13C ~−7‰) would have to have been added to the system. This implies an increase of almost 61 times the present-day volcanic and hydrothermal activity, assuming a yearly flux of mid-oceanic ridge (MOR) CO2 of 0.0082* 1015g (Pytkowicz, 1983) and a maximum time span for the isotopic perturbation of 1* 105 years. It seems that massive release of volcanogenic CO2 may not fully explain the observed magnitude of the negative δ13C excursion. Estimation of the burial of gas hydrate is controversial. Global estimates of hydrate-bound gas in marine sediments vary from the earliest by Trofimuk et al. (1973) to the most recent by Milkov (2004), and are thought to be uncertain. The most widely cited estimate is 10,000 Gt (Kvenvolden, 1988, 1999; MacDonald, 1990; Kvenvolden and Lorenson, 2001) of carbon in clathrates. Milkov (2004) presented a revised global methane hydrate estimate and suggested a maximum 2500 Gt of methane carbon. According to the calculation mentioned above (Spitzy and Degens, 1985), ~2600 (Gt) of methane carbon (δ13C = −65‰ to −80‰) was required to cause the changes in δ13C exhibited in Tethyan successions. This implies a release of ~26% of the widely cited estimate of 10,000 Gt of global hydrates. Therefore, these calculations indicate that a drop in organic activity, rapid oxidation of living biomass, and a massive release of volcanogenic CO2 were insufficient to achieve the observed negative δ13C excursion at F-F boundary. The gas hydrate reservoir is nonetheless further considered an important component of the global carbon cycle. Still another possible cause is oceanic overturn, which has mostly been proposed for the Permo-Triassic ecosystem catas-

trophe. Malkowski et al. (1989), Hoffman et al. (1991), Kimura et al. (1997), and Gruszczy ski et al. (2003) explained the sudden negative excursion of δ13C and extinction events by introducing the idea of a two-box model of the ocean, or two alternative states of the ocean. The two-box or stratified ocean was formed under the warm and ice-free environment (Berry and Wilde, 1978; Keith, 1982; Wilde and Berry, 1984, 1986, and others). The original idea proposed by Degens and Stoffers (1976) considered that the alternation of density-stratified ocean and mixed ocean states could have a significant influence on Earth’s biota. The greenhouse condition in the late Frasnian, documented by various mineralogical and geochemical lines of evidence, is likely to have reduced oceanic circulation, resulting in a stagnant ocean. The triggering mechanism for the oceanic overturn is not clear. The bolide impact (Sandberg et al., 1988, 2002; Geldsetzer et al., 1993) in the ocean was a candidate for mixing a stratified ocean at the end-Permian (Kajiwara et al., 1994), but it is a rather improbable mechanism for the F-F interval (see McGhee, 2001; Reimold et al., 2005). Significance of Negative Excursion of δ18O The variation in δ18O of Late Devonian carbonates is less marked than that of δ13C. The δ18O values of the Late Devonian record are lighter than those of modern carbonates, having values of +1‰ to –4‰ (Milliman, 1974). Cathodoluminescent microscopy of samples from south China indicates that diagenesis is not significant in that area. Then the observed trend toward lighter values of δ18O in the F-F boundary interval of the Liujing section, which is confirmed by data from another section in south China (Wang et al., 1991), together with negative δ18O at the F-F boundary of the Syv`yu River section, can be a sign either of lighter oxygen input to seawater or of seawater warming. We examine the following two aspects of climatic change as explanations for the trend toward lighter δ18O values: 1. Input of light oxygen by the melting of glaciers. The effect of continental glaciation on the isotopic composition of the ocean derives directly from isotopic fractionation in the meteoric cycle (Anderson and Arthur, 1983). The field evidence in the studied areas, including ferruginous beds and mineralogical data such as an increased kaolinite/illite ratio (Gharaie et al. 2004; see also Devleeschouwer et al., 2002), indicates a greenhouse condition at the F-F transition. Consequently there is no any geochemical or field evidence in the studied sections for oceanic cooling. The Late Devonian is generally known as the greenhouse age without glaciers (Frakes et al., 1992; Hallam and Wignall, 1997). A warm and humid climate during the late Frasnian is supported by mineralogical and paleobiological proxies (e.g., Racki, 1998, 1999; Chen et al., 2002; Devleeschouwer et al., 2002; Hladil, 2002) and by numerical modeling (Ormiston and Oglesby, 1995). However, evidence for climate cooling at the F-F boundary was also highlighted (see review in McGhee, 1996). This is largely based on

Chemostratigraphy of Frasnian-Famennian transition study of palynomorphs (Streel et al., 2000), and supported by positive δ18O excursions with high ranges of +1‰ to +1.5‰, derived from biogenic phosphates (Joachimski and Buggisch, 2002). In high-latitude regions, evidence for polar ice caps is recorded only from the uppermost Famennian deposits (Sandberg et al., 1988; Streel et al., 2000). The global paleotemperature curve by Joachimski et al. (2004), derived from worldwide O-isotopic values measured from conodont apatite, strongly suggests that following the greenhouse perturbation spanning the KW crisis, climatic conditions stabilized to a very warm climate (30–33 °C), which prevailed through the early Famennian (beginning with the crepida conodont zone). Chen et al. (2005) also supposed rapid climate changes during the KW crises: an initial increase in the greenhouse effect due to volcanic outgassing, and subsequent climate cooling resulting from decreased atmospheric CO2 levels the burial rate of organic matter increased (see also Buggisch, 1991). These new data contradict the paleobotanical evidence of Streel et al. (2000) suggesting a cold, dry climate, but they are consistent with the data presented in this study. These lines of evidence suggest that supply of light oxygen from deglaciation is not feasible. 2. Higher temperatures of seawater. Although it would not be particularly valuable to obtain a definitive temperature using the δ18O signature of whole-rock samples, the general trend and relative changes in δ18O could be considered an environmental indicator, at least for the samples from south China, which are not strongly altered by diagenesis. Thus the relatively lighter values near the F-F boundary may indicate a higher seawater temperature in this interval than during the rest of Frasnian and Famennian time. Significance of 87Sr/ 86Sr The relatively high 87Sr/86Sr values during the F-F transitional interval compared to those of Frasnian and Famennian age are considered significant (Figs. 3 and 7). The elevated Sr isotope ratio is interpreted as an increase in radiogenic 87Sr influx from old continental rocks into the basin caused by chemical weathering under a warm and humid climate that gradually began earlier in the Frasnian. A global-scale increase in chemical and mechanical erosion of continental rocks, as well as enhanced terrestrial inputs to the oceans, are apparent around the F-F boundary (Averbuch et al., 2005). An incipient uplift is corroborated by the occurrence of a well-dated late Frasnian unconformity within the Moroccan Variscides (Echarfaoui et al., 2002). Major mountain belts were located within equatorial domains (Fig. 5 in Averbuch et al., 2005), and rock uplift and erosion thus occurred in warm and humid climatic conditions characteristic of the late Frasnian. In light of their detailed 87Sr/86Sr data, Chen et al. (2005) interpreted the source of amplified volcanic emanation of CO2, succeeded by accelerated chemical weathering due to the rapid expansion of terrestrial vascular plants in the F-F interval.

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In such conditions, chemical weathering of rocks is likely to have been particularly enhanced (e.g., Francois et al., 1993) and a major increase in oceanic input of the dissolved fraction would be expected. The clay mineral index (Parrish, 1998) is consistent with chemical weathering of rocks: The high abundance of kaolinite relative to illite suggests intensified chemical weathering (Robert and Kennett, 1994, 1997) rather than mechanical erosion. The kaolinite/illite ratio shows striking stratigraphic variation, with ratios increasing in the terminal Frasnian (see Fig. 11 in Gharaie et al., 2004), indicating a climate change to more humid and warmer conditions in the terminal Frasnian. Significance of Redox-Sensitive Elements Under euxinic conditions, such as in the present Black Sea or Cariaco Trench, the trace metal patterns differ from those observed in oxygen-rich waters. Many trace metals undergo a dramatic decrease in solubility at the O2/H2O boundary (e.g., Spencer et al., 1972; Jacobs et al., 1985, 1987; Brumsack, 2006) and are removed from the water column, usually by precipitation as sulfides, and finally buried in the sediments. Elements such as U, As, and V, and (Cu+Mo)/Zn and V/(V+Ni) ratios are indicators of redox of seawater; these exhibit high concentration at the boundary interval in both the Liujing section and the Syv`yu River section (see Figure 7 in Yudina et al., 2002). Parallel laminations in the black shales in Iran strongly support a prevailing euxinic condition during the F-F boundary interval (Gharaie, 2002), as does the occurrence of brown, cubic molds of pyrite pseudomorphs in the Liujing section, which are restricted to the 15-cm-thick micritic limestone of the F-F boundary. Black shale horizons and anomalous concentration of redox-sensitive elements are also reported in other F-F boundaries worldwide (see Geldsetzer et al., 1987; McGhee, 1996; Wang et al., 1996; Joachimski et al., 2001; Ma and Bai, 2002; Yudina et al., 2002; Tribovillard et al., 2004). Fluctuating redox states during the late Frasnian are indeed confirmed in many shelf regions (see Racki et al. 2002; Bond et al., 2004). However, V/(V+Ni) and U/Al ratios are reliable indices of anoxia (Bellanca et al., 1996), and in the Syv`yu succession they suggest fairly uniform, mostly anoxic to euxinic regimes in deep-water settings during the F-F interval (Yudina et al., 2002; see also Tribovillard et al., 2004). Concentration of redox-sensitive elements in the F-F boundary transition in the Liujing section is influenced by the nature of the carbonate (calcite) fraction, which lacks clastics particles. It may be inferred that the high relative abundances of Mo, U, As, Cu and Zn reflect authigenic enrichment resulting from paleoenvironmental conditions; i.e. the anoxic conditions that prevailed across the F-F boundary were also intense in the south China domain (see also Ma and Bai, 2002; Chen et al., 2005). This interpretation can be compared with the results from paleoecological studies. Lethiers et al. (1998) studied benthic or pseudoplanktonic ostracodes and concluded that bottom-water anoxia developed during uppermost Frasnian to lower Famennian times, and that anoxia rose high up into the water column recurrently (see also confirm-

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ing biomarker evidence in Joachimski et al., 2001). By interpreting rare earth element data from conodont phosphate in the La Serre (France) and Kowala Quarry (Poland) sections, Girard and Albaréde (1996) also concluded that there was global anoxia in deep basinal settings. Thus, both the geochemical and paleoecological approaches indicate that the bottom waters in the oceanic environment endured marked anoxic conditions. PROPOSED SCENARIO The δ13C negative excursion, the δ18O, and the anomalous enrichment of redox-sensitive elements are restricted to very narrow stratigraphic range at the F-F boundary. The duration is estimated as much shorter than 1 m.y., probably about several hundred k.y., judging from the decimeter thickness and the assumed rate of sedimentation (Chen and Tucker [2003] suggest ~450 k.y. for the duration of the F-F biotic crisis, based on the time span of a third-order sea-level fall and rise; see also calculation in Sandberg et al., 1988). The short synchronous spikes in abundance of redox-sensitive elements and the related major biotic catastrophic change strongly suggest that the poorly oxygenated event is the cause of the mass extinction event at the F-F boundary, and the δ13C event should have a causal link with the poorly oxygenated condition of seawater (see Hallam and Wignall, 1997; Bond et al., 2004). It is clearly visible that mass extinction followed the onset of the negative shift in the δ13C trend in Paleotethyan successions. Thus, reasons for that shift could be primary causes of the F-F mass extinction. Global warming is inferred to have caused the release of methane, CH4, from clathrates. Knowing the magnitude of the excursion, one can estimate the amount of methane that could have produced it. At least 26% of the total hydrates beneath the seafloor would have to be dissociated to establish such the δ13C excursion exhibited in studied successions. This proportion is derived from the recent reservoir estimation (see above), which is probably an underestimation for the warmer Devonian ocean (see Archer et al., 2004). Another source of relatively enriched in the light carbon isotope such as volcanic CO2 degassing also may explain the negative δ13C excursion in Paleotethyan successions. Volcanism as a Prime Trigger The input of carbon to the atmosphere from volcanic CO2 degassing leads to extremely high levels of atmospheric CO2 and dissolved CO2 in the ocean (Grard et al., 2005). A review of Late Devonian volcanism shows a large igneous province in the peripheries of the Paleotethys Ocean (e.g., Wilson and Lyashkevich, 1996; Racki, 1998, 2005; Gharaie et al., 2004). This province is exemplified by the East European Platform (EEP), which includes one of the world’s largest alkaline massifs, comprising 25 magmatic centers within an area of 100,000 km2 (Kola Province; Wilson and Lyashkevich, 1996); synchronously, associated gabbro-dolerite intrusions crosscut organic-rich strata of Frasnian and earlier age (Wilson and Lyashkevich, 1996). Additional

evidence of rift-related volcanism surrounding the F-F boundary interval has been documented from Siberia and south China (e.g., Racki, 1998, 2005; Veimarn et al., 1998, 2004; Abbott and Isley, 2002; Ma and Bai, 2002; Courtillot and Renne, 2003), and also from north China (Tarim; Hao et al., 2003). This suggests that a cluster of mantle plumes (Wilson and Lyashkevich, 1996; Abbott and Isley, 2002; Courtillot and Renne, 2003; Dobretsov, 2003) could have influenced the thermal and geodynamic evolution of at least the eastern Laurussian-Siberian domain. Volcanic degassing itself is clearly insufficient to explain the F-F changes in δ13C (using the Spitzy and Degens [1985] formula). However, according to a hypothesis by Svensen et al. (2004), intrusion of voluminous mantle-derived melts in carbon-rich sedimentary strata may have caused an explosive release of methane, which would have been transported to the ocean or atmosphere. Organic carbon is partly converted into methane when sedimentary rocks are heated beyond the gas window (100–200 °C; Hunt, 1996). The numerous sills in the Paleotethyan basin were intruded mainly into organic-rich Lower and Middle Devonian mudstones. Melt intruded into organic-rich sedimentary basins may cause considerably higher carbon fluxes into the atmosphere than a similar volume of erupted magma (Svensen et al., 2004). Degassing of one cubic meter of CO2-saturated basaltic melt may release ~3.6 kg of carbon (Caldeira and Rampino, 1991), whereas a melt intruded into organic-rich mudstones may trigger the release of 25–100 kg of carbon per cubic meter of magma (Svensen et al., 2004). Hence, the carbon flux into the atmosphere would easily be more than an order of magnitude higher if the melt intruded into a relatively organic-rich sedimentary sequence rather than being erupted and degassed at the surface (Svensen et al., 2004). Contact heating of organic material and associated methane venting in the EEP Volcanic Province would have the potential to cause the δ13C excursion if the methane was transported to the ocean or atmosphere. We conclude that release of thermogenic methane during the intrusive phase of the EEP Volcanic Province may have caused the extraordinary warming during the F-F transition. Our observations stress the link between large igneous provinces (LIPs) and global climate changes. The climate impact of a LIP would be considerable if the melt were partially or completely emplaced into carbon-rich sedimentary successions. Volcanic basins thus provide a setting for rapid perturbations of the otherwise steady release of carbon from the sedimentary reservoirs. Several other major LIPs temporally correlated with prominent negative carbon isotope anomalies contain extensive subvolcanic intrusive complexes in carbon-rich sedimentary sequences, including the Siberian Traps (250 Ma; the Permo-Triassic boundary), the Karoo Igneous Province (183 Ma; the Early-Middle Jurassic boundary), and the North Atlantic Volcanic Province (55 Ma; start of the Eocene Epoch) (see summary in Svensen et al., 2004). The signature of the Sr isotope ratio strongly suggests that the F-F boundary interval was a time when the chemical weathering was much stronger than during the late Frasnian to late Famennian interval overall, owing to mountain building and the

Chemostratigraphy of Frasnian-Famennian transition pedogenic impact of expanding forests (e.g., Averbuch et al., 2005; Chen et al., 2005). The short-term, extremely warm climatic conditions, which developed during a gentle and long-term greenhouse period from the late Frasnian to the late Famennian (Gharaie et al., 2004), can therefore be called a “super-greenhouse time” (e.g., Larson, 1991; Kerr, 1998). Volcanism was probably the trigger for the overall greenhouse condition (or for a greenhouse setting punctuated with brief cooling episodes; Joachimski et al., 2004; Chen et al., 2005; Racki, 2005). Volcanism would have increased atmospheric CO2 concentration, thereby contributing to the global climatic warming (Wignall, 2001; see also Grard et al., 2005), as suggested for the Kellwasser Crisis by Buggisch (1991) and Becker and House (1994). Frasnian-Famennian Methane Catastrophe Negative δ18O excursions in the south China and subpolar Urals sections probably indicate increased temperature during F-F time span. Synchrony of negative excursions in both δ18O and δ13C may indicate similarities in their genesis. A cluster of rift-related mantle plumes over vast areas of the Paleotethys realm (i.e., EEP, north Iran, Siberia, and south China) would have increased the geothermal gradient and surficial temperature of the basins. Massive dissociation of methane hydrate (both thermogenic and bacteriogenic) is considered as the main cause of negative δ13C excursion as discussed earlier. Paleogeographic position, basin morphology, and features of the hydrological cycle were probably additional factors controlling the carbon and oxygen cycles in the Paleotethys basin (Hoffman et al., 1991). The position of the semi-restricted (to restricted) Paleotethys ocean may have been a significant factor in the clearly exhibited negative δ13C and δ18O excursion in the studied successions. Whether a particular oceanic location is prone to methane eruptions would be determined by paleogeography and seafloor topography, which change on time scales of tens of millions of years, and by the ocean circulation pattern (Ryskin, 2003). The paleogeography of the Middle and Late Devonian may have led to increased organic activity (see Stock, 1990; Poncet, 1990), intensified volcanic-subvolcanic activity (Wilson and Lyashkevich, 1996; Chen et al., 2005), and intrusion of mantle-derived magma into organic-rich sediments. The subsequent massive release of methane would have greatly increased carbon flux into the ocean, and the accumulation of vast amounts of dissolved methane would have the oxygen minimum zone to rise and stagnant anoxic regions to develop (Ryskin, 2003). Although the tectono-volcanic activity was probably long-lasting, compared with the F-F extinction event (ca. 450 k.y.; Chen and Tucker, 2003), the progressively induced warming might have passed a threshold, which then triggered the catastrophic dissociation of gas hydrate. It is therefore the coupling of methane release with water-column anoxia that would have led to the mass extinction at the F-F boundary event. Redox-sensitive elements exhibited in studied sections result from their being concentrated under euxinic conditions at the F-F boundary.

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Large-scale redeposition and brecciation have been widely recognized near or at the F-F transition. Sandberg et al. (1988) postulated even a worldwide breakup of carbonate platform margins at the peak sea-level drop in the earliest Famennian. The most spectacular evidence for such a high-energy event has been recognized in south China by Bai et al. (1994). They causally linked the catastrophic sedimentation with impact-driven megatsunamis strengthened by a global or at least vast regional, stepwise intensification of rifting. In addition, the rapid sea-level fall is explainable by the tectonoeustatic Cathles-Hallam model of sudden sea-level drops during non-glacial periods (see discussion in Racki, 1998). Therefore, all of the disturbances to the seafloor and to sediment stability at this time may have helped trigger the gas hydrate dissociation and catastrophic methane release (see e.g., Bratton, 1999). Methane is likely the main factor influencing the carbon isotopes ratio of seawater during Frasnian-Famennian boundary time span (Matsumoto et al., 2002). A geochemical cycle showing features comparable to the sudden drop in δ13C values has also been documented for the Permo-Triassic boundary (Berner, 2002), the Early Jurassic (Hesselbo et al., 2000; see also review of Weisert, 2000), and the early Eocene (Svensen et al., 2004), all periods associated with mass extinctions. The intensity of mass extinctions varied throughout the Phanerozoic (Sepkoski, 1989) and may have been related to the intensity of the perturbations of the carbon cycle (Gruszczyński et al., 2003). CONCLUSIONS Examination of the negative spike of δ13C and the causal factors for such a signature strongly suggest elimination of all candidates except the dissociation and combustion of methane in setting off intensive volcanic/seismic activity, paired with falling sea level (Sandberg et al., 1988, 2002; Chen and Tucker, 2003, 2004) and disturbed greenhouse climate (Joachimski and Buggisch, 2002; Chen et al., 2005). This hypothesis was previously announced by Bratton (1999), Chen et al. (2002), and Matsumoto et al. (2002). The most plausible reason of negative δ13C excursion in Paleotethys successions was explained by methane release due to magma being intruded into organic-rich sediments; this resulted in partial conversion of organic carbon into methane, as sedimentary rocks were heated. Contact heating of organic material and associated methane venting to the ocean or atmosphere may have caused the extraordinary warming during the F-F transition and the subsequent release of methane from the methane hydrates. Released methane either through contact heating of organic carbon or directly from gas hydrates had the potential to cause the δ13C excursion at F-F boundary interval. The super-greenhouse environment, evidenced by the various signatures described in this paper, would have been caused by the volcanism, but the sharp negative spike of δ13C is in contradiction with the directly released CO2 from volcanic eruptions. Lighter values of δ18O observed in the short duration of the F-F boundary are a likely indicator of higher seawater tem-

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peratures than during the rest of Frasnian and Famennian time. The increase in Sr isotope ratio is interpreted as an influx into the basin due to regional continental uplift, following intense weathering controlled by rapidly accelerated plant-mediated pedogenesis and by the climatic change during the F-F boundary interval (see summary in Averbuch et al., 2005, and Chen et al., 2005). Redox-sensitive elements exhibit high concentration under euxinic conditions at the boundary time span. The dissociation of gas hydrate is the key suspect that can explain all the proposed evidence in this work. During the F-F boundary interval 21% of marine families were exterminated (Sepkoski, 1982) directly by anoxia. The ultimate cause was volcanism and its effect, the dissociation of methane hydrate. ACKNOWLEDGMENTS We are indebted to the Director of Geological Survey of Iran (GSI),. M.T. Koreiie, as well as M. Ghorashi and A. Saeidi (both GSI Tehran) for the logistic and fieldwork support during this study. Funding was provided by the Ministry of Education, Culture and Science of Japan. A. Navab Motlagh (GSI Tehran), J. Taheri (GSI Mashhad), and S. Monibi (Iran Oil Company, Tehran) assisted during fieldwork. J. Komatsubara and R. Matsuda (both of the University of Tokyo) assisted in laboratory work. We are grateful to their advice, organization, and assistance in making this study possible. We also thank P.B. Wignall and M.E. Tucker for in-depth review and constructive comments. REFERENCES CITED Abbott, D.H., and Isley, A.E., 2002, The intensity, occurrence, and duration of superplume events and eras over geological time: Journal of Geodynamics, v. 34, p. 265–307, doi: 10.1016/S0264-3707(02)00024-8. Alavi-Naini, M., 1993, Paleozoic stratigraphy of Iran, in Hushmandzadeh, A., ed., Treatise on the Geology of Iran, v. 5: Tehran, Geological Survey of Iran, p. 1–492. [In Farsi.] Anderson, T.F., and Arthur, M.A., 1983, Stable isotopes of oxygen and carbon and their application to sedimentologic and paleoenvironmental problems, in Arthur, M.A, Anderson, T.F., Kaplan, I., Veizer, J., and Land, L., Stable isotopes in sedimentary geology. SEPM Short Course Notes, No. 10, p. 1–151. Archer, D., Martin, P., Buffett, B., Brovkin, V., Rahmstorf, S., and Ganopolski, A., 2004, The importance of ocean temperature to global biogeochemistry: Earth and Planetary Science Letters, v. 222, p. 333–348, doi: 10.1016/ j.epsl.2004.03.011. Ashouri, A., 1997, The Devonian-Carboniferous boundary in Ozbak-kuh area: Geosciences Scientific Quarterly Journal of Iran, v. 29, p. 46–51. Averbuch, O., Tribovillard, N., Devleeschouwer, X., Riquier, L., Mistiaen, B., and van Vliet-Lanoe, B., 2005, Mountain building-enhanced continental weathering and organic carbon burial as major causes for climatic cooling at the Frasnian-Famennian boundary (c. 376 Ma)?: Terra Nova, v. 17, p. 25–34, doi: 10.1111/j.1365-3121.2004.00580.x. Bai, S.L., Bai, Z.Q., Wang, D.R., Ma, X.P., and Sun, Y.L., 1994, Devonian events and biostratigraphy of South China: Beijing, Peking University Press, 303 p. Becker, R.T., and House, M.R., 1994, Kellwasser Events and goniatite successions in the Devonian of the Montagne Noire with comments on possible causations: Courier Forschungsinstitut Senckenberg, v. 169, p. 45–77. Bellanca, A., Claps, M., Erba, E., Maunitti, D., Neri, R., Premoli Silva, I., and Venezia, F., 1996, Orbitally induced limestone/marlstone rhythms in the Albian-Cenomanian Cismon section (Venetian region, northern Italy):

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Contents

Preface 1. Cenozoic mass extinctions in the deep sea: What perturbs the largest habitat on Earth? Ellen Thomas 2. A major Pliocene coccolithophore turnover: Change in morphological strategy in the photic zone Marie-PietTe Aubry 3. The Paleocene-Eocene Thermal Maximum in Egypt and Jordan: An overview of the planktic foraminiferal record Elisa Guasti and Robert P. Speijer 4. Calcareous nannofossil assemblages and their response to the Paleocene-Eocene Thermal Maximum event at different latitudes: ODP Site 690 and Tethyan sections Eugenia Angori, Gilen Bernaola, and Simonetta Monechi 5. A review of calcareous nannofossil changes during the early Aptian Oceanic Anoxic Event la and the Paleocene-Eocene Thermal Maximum: The influence offertility, temperature, and pC02 Fabrizio Tremolada, Elisabetta Erba, and Timothy J. Bralower

6. Ecosystem perturbation caused by a small Late Cretaceous marine impact, Gulf Coastal Plain, USA David T. King Jr., Lucille W. Petruny, and Thornton L. Neathery 7. Chemostratigraphy of Frasnian-Famennian transition:

Possibility of methane hydrate dissociation leading to mass extinction Mohammad Hossein Mahmudy Gharaie, Ryo Matsumoto, Grzegorz Racki, and Yoshitaka Kakuwa

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