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CBEP 2009 2009 CBEP Wellington, New New Zealand Zealand Wellington,

Extended abstracts from an international conference in Wellington, New Zealand, 12-15 January 2009

Edited by: Erica Crouch, Percy Strong, Chris Hollis.

GNS Science Miscellaneous Series 18

BIBLIOGRAPHIC REFERENCE Crouch, E.M., Strong, C.P., Hollis, C.J., (editors) 2009. Climatic and Biotic Events of the Paleogene (CBEP 2009), extended abstracts from an international conference in Wellington, New Zealand, 12-15 January 2009. GNS Science Miscellaneous Series 18, 163 p.

ORGANISING COMMITTEE GNS Science: Chris Hollis and Liz Kennedy (convenors), James Crampton, Erica Crouch, Hugh Morgans, Ian Raine, Deanne Houghton, Percy Strong Otago University: Ewan Fordyce, Daphne Lee, Gary Wilson Victoria University: Mike Hannah, John Creech Waikato University: Cam Nelson, Ben Andrew Absolutely Organised: Janet Simes, Michelle Vui (conference managers) FORMATTED AND COMPILED BY: Deanne Houghton and Karen Hayes (GNS Science)

© Institute of Geological and Nuclear Sciences Limited, 2008 ISSN 1177-2441 ISBN 978-0-478-19652-8

Climatic and Biotic Events of the Paleogene

CLIMATIC AND BIOTIC EVENTS OF THE PALEOGENE – CBEP 2009 PREFACE The Paleogene greenhouse world is the last time the Earth experienced pronounced global warming and is a potential window into the future climate as reconstructed atmospheric carbon dioxide concentrations were in the range projected to occur over the coming centuries (~1000 parts per million; Zachos et al. 2008). Moreover, the Paleogene records many fascinating episodes in Earth history, including the catastrophic end-Cretaceous asteroid impact, episodes of pronounced global warming, and the stuttering, uneven progression of climate into the modern “icehouse” world. Such issues are the rationale for the “Climatic and Biotic Events of the Paleogene (CBEP)” Conference, a widerranging meeting of Earth and climate scientists that has been held every 2 to 3 years since the first meeting in 1997. CBEP 2009 was held in Wellington, New Zealand, attracting over 130 participants from 20 countries. Over 4 days of oral and poster presentations, CBEP 2009 demonstrated substantial progress in investigating Earth’s past warm climates, including the introduction of new proxies, more common use of multi-proxy approaches, and closer integration of proxy records and modelling. Three major themes that developed during the meeting are highlighted below. First, the Paleogene appears to have been much warmer at many localities, ranging from the tropic to the poles, than previously reconstructed. The new temperature estimates are based on reappraised traditional (leaf physiognomy and foraminiferal δ18O and Mg/Ca) proxies (e.g. Jaramillo this volume), combined with new proxies such as TEX86, MBT/CBT, ‘snake paleothermometry’ and clumped isotopes (e.g. Hollis et al. this volume). While high tropical temperatures are consistent with high greenhouse gas concentrations, the extreme heat at both northern and southern high latitudes during some intervals remains poorly understood. High resolution records, many dated using orbitally tuned-age models, show that Paleogene climate was highly variable. Examples include large-scale floral shifts at the

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Paleocene–Eocene Thermal Maximum (PETM) (see Wing et al. this volume), substantial increases in insect damage with rising temperature (see Currano et al. this volume), and open-ocean salinity perturbations during the Middle Eocene Climatic Optimum (MECO) that all point to a Paleogene climate subject to large scale global climatic fluctuation. Discussions focused on the drivers and feedback of this variability, including whether the major warmings were in or out of phase with orbital forcing and whether all warmings are associated with negative carbon isotope excursions (and vice versa). Finally, simultaneously explaining Early Paleogene warmth and variability is challenging. The most plausible cause of Early Paleogene global warmth is high greenhouse gas concentration. However, the relationship between radiative forcing and greenhouse gas concentration is considered to be logarithmic, so much more greenhouse gas must be added to the atmosphere to cause warming at initially higher concentrations. As shown in several presentations, models have difficulty reproducing the range of climate variability reported from the Paleogene without invoking unreasonably large forcing. Climate models seem to lack key components or feedbacks that may enhance climatic sensitivity to radiative forcing at high greenhouse gas concentrations. CBEP 2009 fieldtrips were popular and wellattended. 50 people participated in a 6 day preconference excursion to the mid and northern South Island and visited Paleogene sequences spanning terrestrial to deep-marine paleoenvironments. A 2 day post-conference trip to Hawkes Bay, including an overnight stay at Rongomaraeroa Marae (Maori meeting house), was attended by 40 people, and concurrent 5 day excursions to the southern South Island and Chatham Island enabled participants to experience a diverse range of sedimentary sequences in remote locations. This extended abstract volume gives a representative view of the research topics discussed at the CBEP 2009 meeting. The papers span a wide range of disciplines,

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techniques, geographic locations and temporal scale, which is part of the motivation for organising and attending this type of multidisciplinary meeting. We now look forward to the next CBEP meeting, to be held at Salzburg, Austria, in 2011. Erica Crouch, Chris Hollis, Matt Huber and Percy Strong.

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REFERENCES Zachos, J.C.; Dickens, G.R.; Zeebe, R.E. 2008: An early Cenozoic perspective on greenhouse warming and carbon-cycle dynamics. Nature 451: 279-283.

Note: parts of this preface have been taken from a meeting review published in EOS (Hollis, C.; Huber, M. 2009: Evolving views on a dynamic greenhouse earth. EOS 90, 22: 194.)

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UNIQUE EOCENE SUBTROPICAL FLORA OF THE RAICHIKHA COAL FIELD (AMUR RIVER REGION, FAR EAST, RUSSIA) M.A. Akhmetiev and V.N. Beniamovski Geological Institute, Russian Academy. Sciences, Pyzhevsky Pereulok, 7, Moscow, 119017, Russia: [email protected]; [email protected]

The Lower Paleogene deposits of Zeya-Bureya sedimentary Basin include two types of Early Paleogene flora: Tsagayan-type (Paleocene warm-temperate) and Raichikha-type (Early Eocene, subtropical). The Paleocene and Eocene floras distributed along the eastern margin of the northwest Pacific from Eastern China and Southern Japan (south) to Kamchatka and Koryakia (north) are distinguished by various conifers, reduced ferns, and diverse warm temperate deciduous taxa representing Myricaceae, Betulaceae, Ulmaceae and the form genus Trochodendroides. Rare evergreen plants and palm Sabal appeared at the Early Eocene global thermal events. The Early Eocene flora of the Raichikha coal field (50°N) is, however, not descended from the Tsagayan Flora and is typically subtropical instead. It contains migrants from the southern region of China which penetrated to the North during Early Eocene warm episode. Modern analogues of the Raichikha flora are widespread through and south China and also Japan (30°N). The Raichikha flora lacks palms, but is characterized by diverse ferns, various evergreen and deciduous conifers. Amentiflorae and Trochodendroides are not represented. Raichikha and Late Tsagayan (predecessor) floras almost completely lack common elements (besides some ferns, Taxodium, Porosia and Platanaceae). Tsagayan-type flora therefore could not be ancestral to the Raichikha flora. It was formed by southern elements that migrated during Early Eocene warming. The plants were preserved in white clay lenses within alluvial-proluvial sands representing oxbox lakes. Kryshtofovich (1946, 1952) described this flora the first time and considered it to be a new link of the Far East Paleogene flora. Baikovskaya devoted attention to the presence of different Filices, Extended Abstracts

Lauraceae and Fabaceae in the flora. Akhmetiev (1973) supposed that the Raichikha flora is one of warmest Eocene Flora of Russian Far East (more than 40% of the leaves have entire margins). Fedotov described more then 100 species, 45% of which have entire margined leaves. He noted that for several tens species it is still necessary understand their systematic position. The Raichikha flora, especially in aquatic and periaquatic assemblages, is similar to the same assemblages of the Paleocene–Early Eocene Zaissan Lake floras (Kiin-Kerish-1), some Eocene Primorie, North Japan and North East China floras (Fushun, coast of Amur gulf near Vladivostok etc.). The Eocene subtropical flora of Japan is characterized by more ecological and morphological similarities than by systematic ones. Aquatic and periaquatic assemblages include: Marchantites spp., Cyclosorus rajczichensis (Fedot.) Akhmet., Osmunda sachalinensis Krysht., Woodwardia sp., Salvinia prearticulata Berry, Regnelidium amurense Fedot., Taxodium dubium (Sternb.) Heer, Nelumbo protospeciosa Sap., Nuphar sp., Hibiscus sp., Limnobiophyllum scutatum (Dawson), Krassil., Zingiberopsis magnifolia (Knowlton) Hickey, Peltandra zaisanica (Fedot.) Fedot., Typha sp., Dicotylophyllum strelkovii Fedot. Some of the aquatic and near-aquatic plants correspond to those known in North American Eocene floras. COMPOSITION AND PHYSIOGNOMY Hepaticopsida–1, Coniferopsida–10, Monocotyledones 6.

Equisetopsida–1, Dicotyledones–71,

Dicotyledones. Taxon with entire margin leaves: 29–44.6%; non entire margin leaves: 36–54.6%; non lobed: 46%.

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The most characteristic taxa of the Raichikha Flora are Anemia elongata (Newb.) Knowlton (Schizaceae), Cyclosorus rajczichensis (Fedot.) Akhmet. (Polypodiaceae), Regnelidium amurense Fedot. (Marsiliaceae), Salvinia preauriculata Berry (Salviniaceae), Taxodium olrikii (Heer) Brown (Taxodiaceae), Neolitsea sp., Lindera rahczichensis Fedotov (Lauraceae), Nelumbo protospeciosa (Sap.) (Nelumbonaceae), Celtis sp. (Ulmaceae), Ficus yubariensis Endo (Moraceae), Urticaceae gen. et sp. indet., Diospyros ficoidea Lesq. (Ebenaceae), Hibiscus sp. (Malvaceae), Mallotus sp. (Euphorbiaceae), Robinia amurensis Baik., and other fruits and leaflets Fabaceae, Myrtophyllum amurensis Fedot. (Myrtaceae), Rhus mixta Kamaeva (Anacardiaceae), Melia sp. (Meliaceae), Ailanthus confucii Ung. (Simaroubaceae), Cardiospermum sp., Delavaya fraxinifolia Fedotov (Sapindaceae), Gouania sp., Zizyphus sp. (Rhamnaceae), Viburnum sp. (Caprifoliaceae), Vitis sp. (Vitaceae), Peltandra zaissanica (Fedot.) (Araceae), and ?Spirodela magna MacGinitie (Lemnaceae). The Raichikha Flora, with a high degree of endemism, lived at the ecotonic zone between evergreen subtropical and mixed mesophytic forests in subtropical seasonal climate.

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CONCLUSION •

The Raichikha Flora originated in the northern part of the Eastern Asia Subtropical realms, near the southern boundary of the Boreal realm.

•

During Raichikha time, the northern boundary of the subtropical zone migrated at least 20°N relative to its present position. The Early Eocene climate was warm temperate to subtropical monsoon (summer wet). This is confirmed for the Raichikha Formation by active kaolinization.

•

The Raichikha flora shares common elements with the Early Eocene flora Zaisan Lake basin, with Northeast China (Fushun) and Japan. Lauraceae and Fabaceae, along with Rhamnaceae (Guoania, Ziziphus), occupied the more elevated habitats, accompanied by ferns such as Lygodium and Anemia. Aquatic and periaquatic plants include rich plant assemblages. ACKNOWLEDGEMENTS

This research was supported by SS-4185.2008.5 and grant No. 08-05-00548 of the Russian Foundation for Basic Research, Programme 15.

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PALEOGENE WESTERN SIBERIAN SEA-STRAIT AND ATLANTIC-ARCTICTETHYS CONNECTIONS OF NORTH-CENTRAL EURASIA: PALEONTOLOGICAL, STRATIGRAPHICAL, SEDIMENTOLOGICAL, PALEOGEOGRAPHICAL AND PALEOCLIMATIC ASPECTS M.A. Akhmetiev, V.N. Beniamovski and T.V. Oreshkina Geological Institute, Russian Academy. Sciences, Pyzhevsky Pereulok, 7, Moscow, 119017, Russia: [email protected]; [email protected]

During the Paleocene–Early Eocene the WestSiberian basin (>2 x 106 km2) was part of an interconnected marine system with the Atlantic, Arctic and Tethys. Atlantic connections existed through straits in the northwestern part of the West Siberian basin: Kara, Novya Zemlya, Severnaya Zemlya gates and Kara-sea strait. The connections with the Arctic basin were through the Severnaya Zemlya gate, Kara-sea strait and also the Khatanga strait. The West Siberian Sea was connected with adjacent Tethyan epicontinental seas (West European and Turanian) through the Orsk and Turgaian gates (Beniamovski 2007). North Atlantic and North Sea-Danish basin water masses and biota moved by circum-North European currents along the northern margin of Europe to reach the West Siberian basin through Kara, Novaya Zemlya gates and Karasea strait. A quasi-upwelling polar stream flowed out from the Paleoarctic Amerasean deepwater cap through Severnaya Zemlya gate and Kara-sea strait along east slope of the Urals, creating diatoms blooms and biogenic siliciceous accumulation in the West Siberian basin. Southern (Peri-Tethys) warm water masses and biota flowed and migrated along the east margin of the West Siberian Sea through Khatanga strait, Kara-sea strait and Severnaya Zemlya gate to the Arctic basin. They were connected with a warm stream from the Tethys through the Turgaian and Orsk gates (Beniamovski 2007; Akhmetiev et al. 2009). In the Late Lutetian, all of the northern part of the West Siberian Sea emerged and transformed into a vast sea-bay and connected with the epicontinental Turanian sea of (northern periphery of Tethys) through the Turgaian sea-strait. Since the Late Lutetian, the biota changed considerably, with the disappearance of siliceous plankton (radiolaria and diatoms) and organic-walled palynomorphs became dominant. This change is reflected in

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the regional biostratigraphical zonal scheme of Akhmetiev and Beniamovski (2004). During the late Middle Eocene–Late Eocene the paleoclimate became semi-arid and subtropical, as opposed to the seasonal monsoon Early Eocene environment (Akhmetiev et al. 2009). Geologically brief migrations from the Peri-Tethyan (Turanian sea) of single Late Lutetian–Early Bartonian planktonic foraminiferal taxa (Pseudogloboquadrina, Subbotina, Catapsydrax, Acarinina, Hantkenina, and Turborotalias) occur in the south-eastern part of West-Siberian sea-bay through the Turgaian sea-strait (Beniamovski 2007). At the Eocene– Oligocene transition the sea retreated from the West Siberian plate and Turgaian depression completely. ACKNOWLEDGEMENTS This research was supported by SS-4185.2008.5 and grant No. 08-05-00548 of the Russian Foundation for Basic Research. REFERENCES Akhmetiev, M.A.; Beniamovski, V.N. 2004: Paleocene and Eocene of western Eurasia (Russian sector)–stratigraphy, paleogeography, climate. Neues Jahrbuch für Geologie und Paläontologie 234: 137-181. Akhmetiev, M.A.; Beniamovski, V.N. 2009: Paleogene floral assemblages around epicontinental seas and straits in Northern Central Eurasia: proxies for climate and paleogeographic evolution. Geologica Acta 7: 297-309. Beniamovski, V.N. 2007. The Paleogene straits of northern Eurasia. In: Baraboshkin, E. Yu. ed. The Cretaceous and Paleogene straits of Northern hemisphere. Geological Faculty of Moscow State University. Pp. 80-118 (in Russian).

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FEATURES OF THE PALEOGENE “WARM” TO “COLD” BIOSPHERE TRANSITION IN THE MIDDLE LATITUDES OF CENTRAL EURASIA (STRATIGRAPHY, PALEOGEOGRAPHY, BIOTA, CLIMATE AND EVENTS) M.A. Akhmetiev, V.N. Beniamovski, N.I. Zaporozhets, T.V. Oreshkina and K.A. Pechnikova Geological Institute, Russian Academy of Sciences, Pyzhevsky Pereulok, 7, Moscow, 119017, Russia: [email protected]; [email protected]

The meridianal marine communicational system Tethys and Arctic Basin persisted until the Lutetian. The climate and biota were subtropical (Ypresian) to transitional subtropical to warm temperate (Lutetian). The narrow-leaved xeromorphic Bartonian flora (Baky, South Ural) formed after closure of the sea-way connecting the Tethys and Arctic basins, allowing fresh water to flow south periodically from West Siberia during the Oligocene. It also changed previous incursions of southern seas from the PeriTethys to West Siberia. Concurrently, there was a radiation of herbivorous mammals and increased endemism of the Asiatic Ungulata. The diversification of brontotheres and tapiridians coincided with the appearance of the meso-xerophyllic flora with evergreen and deciduous elements. At the beginning of the Oligocene, humid and arid phases changed to a general cooling trend, as determined in the authors’ studies of Oligocene mid-latitude sections of Central Eurasia. These exhibited a rhythmic intercalation of layers with palynological spectra dominants: Pinus, xerophytic Quercus, Ephedra, Chenopodiaceae, Ericaceae pollen, reduction spores (xerothermic phases) with strata containing dominant mesophytic and hygrophytic taxa palynomorphs: ferns, Taxodiaceae and reduction herbaceous pollen

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(humid phases). At least five such inversions occurred in the Early Oligocene. Later, different plants of this Flora appeared as elements of the Oligocene Turgai Flora. At the Eocene–Oligocene transition there were two types of Flora: 1) a deciduous, mesophyllic Turgaian type Flora representing riparian facies; and 2) a small-leaved, xeromorphic Flora with Lauraceae, and evergreen Fagaceae, distributed on the higher plains and hills. Gigantic Indricotheridae occupied the region during the mid Oligocene. The climate became more arid and open areas appeared. The Turtass “lake-sea” containing an unusual biota formed at the Early–Late Oligocene transition, the results of surface runoff ponding during one of the last ingression of the Turgai strait. It was several time larger then recent large lakes. Water of the Turtass “lake-sea” drained to the north during the Oligocene– Miocene Transition. In the Late Oligocene the climate became more humid and the Turgaiantype deciduous mesophyllic Flora spread across vast areas. ACKNOWLEDGEMENTS This research was supported by SS-4185.2008.5 and grant No. 08-05-00548 of the Russian Foundation for Basic Research.

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TURNOVER AT THE SEA FLOOR DURING THE PALEOCENE–EOCENE THERMAL MAXIMUM: EVIDENCE FROM THE WESTERN TETHYS Laia Alegret1 and Silvia Ortiz1,2 1

Departamento de Ciencias de la Tierra, Universidad de Zaragoza. 50009 Zaragoza, Spain; 2Department of Earth Sciences, University College London, WC1E 6BT London, UK.

INTRODUCTION The greenhouse world of the Paleogene underwent significant disruption at the Paleocene–Eocene transition, when temperatures rapidly increased by 5 to 9ºC in the oceans and on land. During this episode, which is commonly known as the Paleocene– Eocene Thermal Maximum (PETM), a major perturbation of the global carbon cycle occurred, including a negative carbon isotope excursion (CIE) in marine and terrestrial δ13C values of carbonate and organic carbon and a ~2 km shoaling of the calcite compensation depth in the deep sea (Kennett and Stott 1991; Zachos et al. 2005). This global perturbation of the carbon cycle is interpreted in terms of a rapid input of isotopically light carbon into the ocean-atmosphere system, possibly related to the massive dissociation of marine methane hydrates, although the triggering mechanism is still under debate. The onset of the PETM was characterized by abrupt changes such as the acidification of the oceans and rapid changes in terrestrial and marine biota, including the largest extinction of deep-sea benthic foraminifera (30–50% of the species) recorded during the Cenozoic, a rapid evolutionary turnover of planktic foraminifera and calcareous nannoplankton, the global acme of the dinoflagellate genus Apectodinium and its migration to high latitudes, the rapid diversification of shallowwater larger benthic foraminifera, latitudinal migration of plants, and a rapid radiation of mammals on land (see references in Thomas 2007). Although the biotic response to the PETM has been studied intensively, the linkages between carbon cycle perturbation and coeval biotic changes are not completely understood. Individual ecosystems may have responded directly to such aspects of environmental change as carbon addition (ocean acidification, elevated pCO2) or indirectly to consequences of carbon release, such as rising temperatures, increased

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precipitation, and changes in nutrient supply or distribution (Bowen et al. 2006). The study of the benthic foraminiferal extinction event (BEE) is particularly interesting because deep-sea benthic foraminifera had survived without significant extinction through global environmental crises such as that related to the asteroid impact at the end of the Cretaceous (e.g. Alegret and Thomas 2005). The BEE lasted ~10 k.y. or less (Thomas 2007; Alegret et al. 2009a, b), and its cause (oxygen deficiency at the seafloor, increased corrosivity of the waters for CaCO3, changes in productivity, etc.) is not clear. For example, data from open ocean sites do not support global hypoxia, and are inconsistent with regard to global productivity changes (e.g. Thomas 2007). More detailed analyses of benthic foraminifera across the PETM are thus needed to look into the causes of the BEE. However, very few locations are available where the sequence of events at the onset of the PETM can be studied, mainly because of severe carbonate dissolution, which often causes incompleteness of records (Zachos et al. 2005). The sharp decrease in carbonate content or carbonate dissolution also makes it difficult to correlate the PETM marine record, due to problematic estimations of the time involved between the onset of the CIE and the lowest carbon isotope value (Petrizzio 2007). Several studies on the microfossil turnover at the bathyal-abyssal Alamedilla section (Subbetic Cordillera; southern Spain) have pointed out the completeness of the sedimentary record across the PETM (Arenillas and Molina 1996; Monechi et al. 2000; Alegret et al. 2009b), especially because the base of the CIE occurs in a short interval with a high calcite content (Lu et al. 1996). THE PETM RECORD AT ALAMEDILLA A continuous succession of upper Paleocene and lower Eocene pelagic sediments is very 9

Atlantic Ocean


Zumaya

Contessa

Caravaca Alamedilla

Tethys Dababiya GSSP P/E

Figure 1 Paleogeographical reconstruction of the studied area 55.8 million years ago, and location of the Alamedilla section and other sites and sections referred to in the text.

well exposed at Alamedilla section, about 2 km south of the village of Alamedilla (Granada province, southern Spain; Fig. 1). Geologically, this section is located in the central Subbetic Zone of the Betic Cordillera. Upper Paleocene sediments consist of gray marls, with a 15 cm turbiditic layer intercalated in the lower part of the studied section (Figs 2 & 3). In contrast, the lowermost Eocene is represented by a distinctive, 30 cm red clay interval that is overlain by red marls and, higher up in the section, by gray marls (Fig. 2). The onset of the carbon isotope excursion (CIE) was identified by Lu et al. (1996, 1998) in sample 13.25, in a level of pink marls that grade into the red clay interval (13.50–13.80 m; Figs. 2 & 3). According to the definition of the Global Stratotype Section and Point (GSSP) for the P/E boundary at Dababiya, Egypt (Aubry et al. 2007), the boundary must be placed in sample 13.25 at Alamedilla, coinciding with the onset of the CIE. Planktic and benthic foraminiferal δ13C values show a negative excursion of 1.7‰ in both surface and bottom waters, and δ18O values point to a 4ºC warming in bottom waters, with little change in surface waters (Lu et al. 1998). Coeval with the onset of the CIE, Lu et al. (1996) documented a decrease in the percentage of calcite, whereas the percentage of silicate minerals increased up to 12% of whole rock composition during the core of the CIE.

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For biostratigraphic control, we follow the planktic foraminiferal and calcareous nannofossil zones identified by Arenillas and Molina (1996), Molina et al. (1999) and Monechi et al. (2000) at Alamedilla (Fig. 2). According to these authors, the studied section is biostratigraphically complete. BENTHIC FORAMINIFERAL TURNOVER Changes in benthic foraminiferal assemblages across the PETM at Alamedilla are here described based on the data set published by Alegret et al. (2009b). The size fraction >63 μm was chosen to avoid the loss of small species after the extinction event. The composition of the assemblages across the PETM is shown in Fig. 2. Upper Paleocene, pre-extinction faunas are diverse and heterogeneous, and are dominated by calcareous taxa (Fig. 2). It must be noted that the elongate, small buliminid species Siphogenerinoides brevispinosa has been identified in the uppermost Paleocene at Alamedilla. This species has been reported commonly from PETM Ocean Drilling Program (ODP) Sites, but only from one onland PETM section (Contessa Road section, Italy; Giusberti et al. 2009). This species is very abundant (>30% of the assemblages) in the uppermost Paleocene and lowermost Eocene at Contessa (Giusberti et al. 2009), whereas it makes up to 3% of the assemblages Extended Abstracts


levels (below 1.5 mg O2/L; e.g. Morigi et al. 2001). We calculated the BFOI in our samples (Fig. 2). The index shows overall mean values ~50, with the highest figures observed in the lowermost Eocene, resulting from the high percentages of Globocassidulina subglobosa (an oxic indicator in the modern oceans). In our opinion these fluctuations in the BFOI do not indicate changes in oxygen conditions, but are the result of changes in benthic foraminiferal assemblages, which were

at Alamedilla, and was not identified by Ortiz (1995) in the nearby Caravaca section.

EOCENE A. sibaiyaensis NP 10

P. wilcox.

Scale (m) Samples

(1) (2)

Lithology

Biozones

Age

The BEE is recorded at Alamedilla in coincidence with the onset of the CIE; the BEE spans a 30 cm interval in the lowermost Eocene, and it affected a total of 26 species; the lowermost 20 cm of the BEE are recorded over an interval with a high calcite content (Alegret et al. 2009b), suggesting that ocean acidification was not the main trigger of the 13

δ C

(Lu et al. 1996) 0,5 1

2

BFOI

%CaCO3

(Lu et al. 1996) 30

50

100 30

100

High oxic Nuttallides truempyi, Osangularia spp., Anomalinoides spp., buliminids, Pleurostomella spp., Stilostomella spp.

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15

BEE

A. ber.

50

Low oxic

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Globocassidulina subglobosa and other OPPORTUNISTIC species (Aragonia aragonensis, Tappanina selmensis Reussella terquemi, Abyssamina quadrata)

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13

PALEOCENE M. gracilis NP 9

(Benthic Foraminiferal BENTHIC FORAMINIFERAL ASSEMBLAGES Oxygenation Index)

Onset of CIE=PEB

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Red clay Turbidite Marls

Glomospira acme BEE: Extinction of 26 species

PRE-EXTINCTION FAUNA: Anomalinoides spp., buliminids, polymorphinids, Cibicidoides hyphalus, C. pseudoperlucidus, Nuttallides truempyi, Osangularia, Stensioeina beccariiformis, Siphogenerinoides brevispinosa, Stilostomella spp.

Figure 2 δ13C isotopes in planktic foraminifera and percentages of CaCO3 (modified from Lu et al. 1996, 1998), and BFOI (Kaiho 1994) and composition of the benthic foraminiferal assemblages across the PETM at Alamedilla. (1) Arenillas and Molina (1996) and Molina et al. (1999); (2) Monechi et al. (2000). M.= Morozovella; A. ber.= Acarinina berggreni; P. wilcox.= Pseudohastigerina wilcoxensis; BEE (benthic foraminiferal extinction event; Alegret et al. 2009b); CIE= Carbon Isotope Excursion; PEB= Paleocene/Eocene boundary.

extinctions. In order to assess the oxygen deficiency as a possible cause of the extinctions, we considered the reddish colour of the sediments across the CIE (which indicate oxic conditions; Fig. 3), and the Benthic Foraminiferal Oxygenation Index (BFOI; Kaiho 1994). This author suggested that an oxygenation index could be derived from the benthic foraminiferal faunal composition of calcareous taxa; however, some authors have suggested that the BFOI is not applicable, possibly with the exception of environments with extremely low oxygenation

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triggered by other factors than oxygenation at the seafloor. Assemblages just above the BEE, in samples with the lowest calcite content (13.50 and 13.60), are strongly dominated by agglutinated foraminifera, mainly Glomospira charoides and trochamminids. This assemblage corresponds to the well-known Early Eocene Glomospira Acme, which has been identified in deep-water settings of western Tethys and North Atlantic (e.g. Kaminski and Gradstein 2005).

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ACKNOWLEDGEMENTS LA acknowledges support from a “Ramón y Cajal” research contract from the Spanish Ministry of Science and Technology and the European Social Fund, and SO from a postdoctoral grant from the Spanish Ministry of Science and Technology. This research was funded by project Consolider CGL 200763724. Figure 3 Photograph showing the P-E transition at Alamedilla. Older to younger sediments are observed from right to left. The uppermost Paleocene gray marls and a turbidite level intercalated are seen on the right, and the lowermost Eocene sediments, including the distinctive red clay interval, are observed towards the left.

After careful comparison with the samples immediately underneath the Glomospira peak, Alegret et al. (2009c) recently concluded that this peak was mainly related to; (1) CaCO3 dissolution triggered by the shoaling of the CCD; (2) Dilution of the carbonate compounds by silicicate minerals (as inferred from the increased sedimentation rates); and (3) An ecological component. The later may be related to methane dissociation from gas hydrates or from organic matter heated up by igneous intrusions in the North Atlantic; or to regions associated with volcanic ash deposits from the North Atlantic Volcanic Province. However, further studies are needed to determine whether the Glomospira peak was linked to any of these sources of isotopically light carbon. The abundance of calcareous taxa rapidly recovers ~30 cm above the BEE, where a series of peaks in the abundance of opportunistic taxa such as Globocassidulina subglobosa, Aragonia aragonensis, Abyssamina quadrata, Reussella terquemi or Tappanina selmensis, have been recorded, coinciding with minimum δ13C values. These peaks have been observed across the PETM in other on-land sections (e.g. Alegret et al. 2009a; Giusberti et al. 2009) as well as in ODP and DSDP sites (e.g. Thomas 2007), and their detailed study and chronological calibration may help to understand the effects of the PETM.

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REFERENCES Alegret, L.; Thomas, E. 2005: Cretaceous/Paleogene boundary bathyal paleoenvironments in the central North Pacific (DSDP Site 465), the Northwestern Atlantic (ODP Site 1049), the Gulf of Mexico and the Tethys: The benthic foraminiferal record. Palaeogeography, Palaeoclimatology, Palaeoecology 224: 53-82. Alegret, L.; Ortiz, S.; Orue-Etxebarria, X.; Bernaola, G.; Baceta, J.I.; Monechi, S.; Apellaniz, E.; Pujalte, V. 2009a: The Paleocene-Eocene Thermal Maximum: new data from the microfossil turnover at Zumaia section. Palaios 24: 318-328. Alegret, L.; Ortiz S.; Molina, E. 2009b: Extinction and recovery of benthic foraminifera across the Paleocene-Eocene Thermal Maximum at the Alamedilla section (Southern Spain). Palaeogeography, Palaeoclimatology, Palaeoecology, 279: 186-200. Alegret, L.; Ortiz S.; Arenillas, I.; Molina, E. 2009c: What happens when the ocean is overheated? The foraminiferal response across the Paleocene-Eocene Thermal Maximum at the Alamedilla section (Spain), accepted. Geological Society of America Bulletin. Arenillas, I.; Molina, E. 1996: Bioestratigrafía y evolución de las asociaciones de foraminíferos planctónicos del tránsito Paleoceno-Eoceno en Alamedilla (Cordilleras Béticas). Revista Española de Micropaleontología 28: 75-96. Aubry, M-P.; Ouda, K.; Dupuis, C.; Berggren, W.A.; Van Couvering, J.A. 2007: The Global Standard Stratotype-section and Point (GSSP) for the base of the Eocene Series in the Dababiya sction (Egypt). Episodes 30: 271286. Bowen, G.J.; Clyde, W.C.; Koch, P.L.; Ting, S.; Alroy, J.; Tsubamoto, T.; Wang, Y. 2006: Mammalian Dispersal at the Paleocene/Eocene Boundary. Science 15: 2062-2065. Giusberti, L.; Coccioni, R.; Sprovieri, M.; Tateo, F. 2009: Perturbation at the sea floor during the

Extended Abstracts


Paleocene-Eocene thermal maximum: evidence from benthic foraminifera at Contessa Road, Italy. Marine Micropaleontology 70: 102-119. Kaiho, K. 1994: Benthic foraminiferal dissolved oxygen index and dissolved oxygen levels in the modern ocean. Geology 22: 719-722. Kaminski, M.A.; Gradstein, F. 2005: Atlas of Paleogene Cosmopolitan Deep-Water Agglutinated Foraminifera. Krakow, Poland, Grzybowski Foundation Special Publication 10. 548 p. Kennett, J.P.; Stott, L.D. 1991: Abrupt deep-sea warming, palaeoceanographic changes and benthic extinctions at the end of the Palaeocene. Nature 353: 225-229. Lu, G.; Keller, G.; Adatte, T.; Ortiz, N.; Molina, E. 1996: Long-term (105) or short-term (103) δ13C excursion near the Palaeocene-Eocene transition: evidence from the Tethys. Terra Nova 8: 347-355. Lu, G.; Adatte, T.; Keller, G.; Ortiz, N. 1998: Abrupt climatic, oceanographic and ecologic changes near the Paleocene-Eocene transition in the deep Tethys basin: the Alamedilla section, southern Spain. Ecoglae geologica Helvetiae 91: 293-306. Molina, E.; Arenillas, I.; Pardo, A. 1999: High resolution planktic foraminiferal biostratigraphy and correlation across the Palaeocene/Eocene boundary in the Tethys. Bulletin de la Société géologique de France 170: 521-530.

Extended Abstracts

Monechi, S.; Angori, E.; Von Salis, K. 2000: Calcareous nannofossil turnover around the Paleocene/Eocene transition at Alamedilla (southern Spain): Bulletin de la Société géologique de France 171: 477-489. Morigi, C.; Jorissen, F.J.; Gervais, A.; Guichard, S.; Borsetti, A.M. 2001: Benthic foraminiferal fanunas in surface sediments off NW Africa; relationship with organic flux to the ocean floor. Journal of Foraminiferal Research 31: 350-368. Ortiz, N. 1995: Differential patterns of benthic foraminiferal extinctions near the Paleocene/Eocene boundary in the North Atlantic and the western Tethys. Marine Micropaleontology 26: 341-359 Petrizzo, M.R. 2007: The onset of the PaleoceneEocene Thermal Maximum (PETM) at Sites 1209 and 1210 (Shatsky Rise, Pacific Ocean) as recorded by planktonic foraminifera. Marine Micropaleontology 63: 187-200. Thomas, E. 2007: Cenozoic mass extinctions in the deep sea; what disturbs the largest habitat on Earth? In Monechi, S.; Coccioni, R.; Rampino, M. ed. Large Ecosystem Perturbations: Causes and Consequences. Boulder, Colorado, Geological Society of America Special Paper 424. Pp. 1-24. Zachos, J.C.; Röhl, U.; Schellenberg, S.A.; Sluijs, A.; Hodell, D.A.; Kelly, D.C.; Thomas, E.; Nicolo, M.; Raffi, I.; Lourens, L.J.; McCarren, H.; Kroon, D. 2005: Rapid Acidification of the Ocean During the Paleocene-Eocene Thermal Maximum. Science 308: 1611-1615.

13


THE PACIFIC OCEAN: NO STRANGELOVE OCEAN AFTER THE CRETACEOUS–PALEOGENE IMPACT Laia Alegret1 and Ellen Thomas2,3 1

Departamento de Ciencias de la Tierra, Universidad de Zaragoza, 50009 Zaragoza, Spain; 2Center for Study of Global Change, Yale University, CT 06520-8109 New Haven, USA; 3Department of Earth and Environmental Sciences, Wesleyan University, CT 06459-0139 Middletown, USA.

INTRODUCTION Evidence that a large asteroid impacted Earth at the Cretaceous–Paleogene (K–Pg) boundary (widely known as the K–T boundary) is convincing, but the ecological effects of such an impact and its effects on oceanic ecosystems, thus on the global carbon cycle, are strongly debated. The K–Pg extinction was one of the largest mass extinctions of the Phanerozoic, but deep-sea benthic foraminifera, one of the few groups of organisms that provide information on deepsea benthic ecosystems (the largest habitat on Earth) do not show significant extinction, only transient changes in community structure (i.e. relative abundance and diversity; e.g. Coccioni and Galeotti 1994; Alegret and Thomas 2005; Culver 2003; Alegret 2007; Thomas 2007). These transient changes have been explained by the prevalence of oligotrophic conditions on the seafloor, either as the result of collapse of the primary producers in the pelagic food web (“Strangelove Ocean” model; Hsü and McKenzie 1985) or of the collapse of transport of organic matter to the seafloor due to extinction of pellet-producing zooplankton (“Living Ocean” model; d’Hondt et al. 1998), as recognized in the collapse of carbon isotopic gradients between surface and bottom waters. Detailed studies of the biotic and paleoenvironmental response to the impact may elucidate the mechanisms of oceanic primary productivity and the delivery of organic matter to the seafloor (the biological pump). The pattern of benthic faunal turnover is not clear around the Gulf of Mexico and the North Atlantic because of the proximity to the impact site on the northern Yucatan peninsula (Fig. 1): at these locations the sedimentary record is incomplete or severely disturbed due to the destabilization of the continental margins and mass 14

wasting processes (e.g. Alegret et al. 2001). We thus must study locations distal from the impact site in order to obtain high resolution records of biotic reaction to the impact. Sites in the Pacific Ocean are distal, and represent environmental conditions in the world’s largest ocean basin. We analyzed changes in benthic foraminiferal assemblages at Ocean Drilling Program (ODP) Site 1210 on Shatsky Rise (northwest Pacific), and compared these to the record at Deep Sea Drilling Project (DSDP) Site 465, on Hess Rise (central north Pacific; Alegret and Thomas 2005). SEDIMENTARY RECORD OF THE K/PG TRANSITION IN THE PACIFIC DSDP Site 465 has been above the local Carbonate Compensation Depth (CCD) since its formation, and benthic and planktic foraminifera are well preserved in homogeneous calcareous nannofossil and foraminiferal oozes. The record from DSDP Hole 465A appears to be biostratigraphically complete, but rotary drilling disturbance caused irregular mixing of uppermost Maastrichtian and lowermost Danian sediments in a 20–30 cm zone across the K–Pg boundary. Deposition was at lower bathyal

Figure 1 Paleogeographical reconstruction of the world during the Cretaceous-Paleogene transition, and location of the studied sites and other relevant sites. Extended Abstracts


depths (~1500 m), with diverse benthic foraminiferal assemblages dominated by calcareous taxa of the Velasco-type faunas (Alegret and Thomas 2005). K–Pg boundary sections recovered at ODP Site 1210 at Shatsky Rise (Leg 198, west central Pacific) did not suffer drilling disturbance and represent some of the best preserved deep-sea records of the K–Pg extinction and its aftermath. The sediments include uppermost Maastrichtian white to pale orange nannofossil ooze and lowermost Paleocene, grayish-orange foraminiferal ooze (10 cm), grading upward into a white foraminiferal nannofossil chalk (20 cm), then into a grayish-orange nannofossil ooze. There is intense bioturbation, with up to 5 cm long burrows across the boundary, but the substantial thickness of the uppermost Maastrichtian Micula prinsii (CC26) Zone and the lowermost Danian Parvularugoglobigerina eugubina (P) Zones indicates that the K–Pg boundary is paleontologically complete. We used the orbitally tuned age model for Site 1210 (Westerhold et al. 2008), and correlated this record using carbon isotope data to the record for Hole 465, thus improving on the low resolution age model for that site, and recalculating benthic foraminiferal accumulation rates. MICROFOSSIL TURNOVER AND FOOD PROXIES The well-known mass extinction of calcareous primary producers at the K–Pg boundary has

been documented at Site 1210 (Bown 2005), with a catastrophic mass extinction of calcareous nannoplankton followed by a recovery interval with abundant Cretaceous survivors such as calcispheres (cysts of calcareous dinoflagellates), and successive acmes of Danian taxa, representing rapid colonization by former geographically restricted forms. The benthic foraminiferal assemblages at Site 1210 are diverse and dominated by calcareous taxa (87–100%), with many common species of the Velasco-type fauna, typical for lower-bathyal to abyssal settings (~1500–2000 m). The benthic assemblages, in contrast to the nannoplankton assemblages but as also documented elsewhere (Thomas and Alegret 2005; Alegret 2007, Thomas 2007), did not suffer significant extinction at Site 1210. Despite the difference in paleodepth between Sites 465 and 1210, and the differences in most common taxa at both sites, there are strong similarities between the records across the K–Pg. At both sites, diversity and heterogeneity decreased abruptly across the boundary (Fig. 2), while benthic foraminiferal accumulation rates (BFARs) peaked for about 15 k.y., and there was a somewhat shorter-lived peak in percentage buliminids. The BFAR values peaked at both Sites 465 and 1210 for a similar time period (Fig. 2), while the diversity (expressed as alpha-diversity) was low. The BFAR values were in general higher for the somewhat shallower Site 465, as expected, but with the largest difference during the recovery period. Both the relative abundance of buliminids (an

Figure 2 Diversity (alpha), relative abundance of bulimind taxa and benthic foraminiferal accumulation rates across the K/Pg boundary at Pacific Ocean Sites 465 (Hess Rise) and 1210 (Shatsky Rise). Extended Abstracts

15


indicator for generally high and stable food supply) and the BFAR (an indicator of high food supply to the sea floor) (Gooday 2003; Jorissen et al. 2007) thus peaked directly after the K–Pg extinction (as observed in calcareous nannoplankton and planktic foraminifers) at two sites in the Pacific Ocean, while there was no extinction of benthic foraminifera significantly different from background extinction levels. This observation is in strong conflict with the hypothesis that primary productivity collapsed for several hundred thousand of years, leading to the existence of an extreme oligotrophic, “Strangelove Ocean”, as indicated by the collapse of the carbon isotope gradient between surface and deep ocean (e.g. Hsü and MacKenzie 1985). The observation is likewise in contrast with the “Living Ocean Model” (d´Hondt et al. 1998), which argues for rapid recovery of oceanic primary productivity in terms of biomass (but not in terms of diversity). In this model, however, the organic matter does not reach the benthic foraminifera on the seafloor for several hundred thousand years due to the collapse of the biological pump, i.e. transport of the food to the seafloor. It is extremely improbable that the minor and transient changes in diversity and species composition of benthic faunal assemblages could possibly be the response of the fauna to a major, long-term collapse of oceanic productivity, because in the presentday ocean benthic foraminiferal assemblages are strongly coupled to productivity in the surface waters (Gooday 2003; Jorissen et al. 2007). Benthic foraminiferal assemblages and BFARs point, in contrast, to a high food supply to the seafloor, thus abundant primary productivity as well as a functioning biological pump, during the earliest Danian at both Sites 465 and 1210, i.e. probably over a large part of the Pacific Ocean. We suggest that high-food taxa (e.g. buliminids) may have bloomed at least at some locations, while the vacated niches of primary producers (calcareous nannoplankton) were rapidly filled by others, leading to reorganization of the planktic ecosystem, but not to a global, severe decrease in delivery of food to the seafloor. Such high productivity following the K–Pg extinction can be seen as resulting from the rapid proliferation of “bloom-taxa”, opportunistic species filling the niches emptied by the mass extinction. The

16

high abundances of small biserial and triserial planktic foraminifera at many locations can also be interpreted as reflecting high productivity (e.g. Koutsoukos 1996). The collapsed vertical gradient in carbon isotopes thus remains to be explained. ACKNOWLEDGEMENTS LA acknowledges support from a “Ramon y Cajal” research grant from the Spanish Ministry of Science and Technology and the European Social Fund. This research was funded by project Consolider CGL 200763724. ET acknowledges funding by NSF Grant OCE 720049. REFERENCES Alegret, L. 2007: Recovery of the deep-sea floor after the Cretaceous/Paleogene boundary event: the benthic foraminiferal record in the BasqueCantabrian basin and in South-eastern Spain. Palaeogeography, Palaeoclimatology, Palaeoecology 255: 181-194. Alegret, L.; Thomas, E. 2005: Cretaceous/Paleogene boundary bathyal paleoenvironments in the central North Pacific (DSDP Site 465), the Northwestern Atlantic (ODP Site 1049), the Gulf of Mexico and the Tethys: The benthic foraminiferal record. Palaeogeography, Palaeoclimatology, Palaeoecology 224: 53-82. Alegret, L.; Molina, E.; Thomas, E. 2001: Benthic foraminifera at the Cretaceous-Tertiary boundary around the Gulf of Mexico. Geology 29: 891-894. Bown, P. 2005: Selective calcarous nannoplankton survivorship at the Cretaceous-Tertiary boundary. Geology 33: 653-656. Coccioni, R.; Galeotti, S. 1994: K-T boundary extinction: geologically instantaneous or gradual event? Evidence from deep-sea benthic foraminifera. Geology 22: 779-782. Culver, S. J. 2003: Benthic foraminifera across the Cretaceous-Tertiary (K-T) boundary: a review. Marine Micropaleontology 47: 177-226. d´Hondt, S.; Donaghay, P.; Zachos, J.C.; Luttenberg, D.; Lindinger, M. 1998: Organic carbon fluxes and ecological recovery from the Cretaceous-Tertiary mass extinction. Science 282: 276-279. Gooday, A.J. 2003: Benthic foraminifera (Protista) as tools in deep-water palaeoceanography: environmental influences on faunal

Extended Abstracts


characteristics. Advances in Marine Biology 46: 1-90. Hsü, K.J.; McKenzie, J. 1985: A “Strangelove Ocean” in the earliest Tertiary. Geophysical Monographs 32: 487-492. Jorissen, F.J.; Stigter, H.C.; Widmark, J.G.V. 1995: A conceptual model explaining benthic foraminiferal microhabitats. Marine Micropaleontology 26: 3-15. Jorissen, F.J.; Fontanier, C.; Thomas, E. 2007: Paleoceanographical proxies based on deep-sea benthic foraminiferal assemblage characteristics. In Hillaire-Marcel, C.; de Vernal, A. ed. Proxies in Late Cenozoic Paleoceanography: Pt. 2: Biological tracers and biomarkers. Elsevier, Pp. 263-326.

Extended Abstracts

Koutsoukos, E.A.M. 1996: Phenotypic experiments into new pelagic niches in early Danian planktonic foraminifera: after-math of the K/T boundary event. In M.B. Hart, ed. Biotic Recovery from Mass Extinction Events. Geological Society Special Publication 102: 319–335. Thomas, E. 2007: Cenozoic mass extinctions in the deep sea; what disturbs the largest habitat on Earth? In Monechi, S.; Coccioni, R.; Rampino, M. ed. Large Ecosystem Perturbations: Causes and Consequences. Geological Society of America Special Paper 424: 1-24. Westerhold, T.; Roehl, U.; Raffi, I.; Fornaciari, E.; Monechi, S.; Reale, V.; Bowles, J.; Evans, H.F. 2008: Astronomical calibration of Paleocene time. Palaeogeography, Palaeoclimatology, Palaeoecology 257: 377-403.

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A MIDDLE EOCENE SOUTHERN PACIFIC CLIMATE RECORD FROM HAMPDEN BEACH: THE MECO IN NEW ZEALAND? Catherine E. Burgess1, Paul N. Pearson1, Caroline H. Lear1 and Hugh E. G. Morgans2 1

School of Earth and Ocean Sciences, Cardiff University, Park Place, Cardiff, CF10 3YE, UK: [email protected]; 2GNS Science, P.O. Box 30368, Lower Hutt, New Zealand.

The Hampden Formation on the east coast of South Island, New Zealand (Fig. 1) is a calcareous silt to very fine sandstone deposited in a shelf edge environment (Fig. 2) containing excellently well-preserved carbonate microand nannofossils (Morgans in press; Pearson and Burgess 2008). The geochemistry and abundance records of these microfossils, combined with the sedimentology of the site, provide an important record of Middle Eocene Paleoclimate (Burgess et al. 2008). The oxygen isotopic composition (δ18O) of well-preserved carbonate microfossils records warmer temperatures than previously estimated for 55°S (Shackleton and Kennett 1975; Zachos et al. 1994), strongly suggesting that during the middle Eocene New Zealand lay in the path of a relatively warm southward flowing current (e.g. Nelson and Cooke 2001) rather than an Antarctic gyre (e.g. Huber et al. 2004). The δ18O record through the Hampden Formation shows a trend of decreasing sea surface temperature from an average of ca 18°C at the base of the section (ca 42 Ma) to an average of 14°C the top at (ca 39 Ma). The formation also shows sedimentary evidence for decreased chemical weathering through this

Figure 1 cyclicity.

18

period, more likely due to decreased temperatures or increased runoff than increased ice. The section records a transient warm interval between ca 40.0 and 39.6 Ma — this excursion is ca 450 ka in duration, exhibiting a δ18O decrease of ca 0.5 ‰, and corresponding to a maximum temperature increase of ca 2.5°C. During this warm interval in the marine record there is an incursion of Hantkenina australis, a tropical foraminfer species and an increase in percentage abundance of foraminifera belonging to ‘mixed-layer’ species. There is also an increase in the ratio of chemical to physical weathering in the terrigenous source area seen as a dramatic decrease in the weight % >63 µm of the sediment, largely due to an increase in the ratio of clays to feldspars. This warm excursion is considered to be the first New Zealand record of the Middle Eocene Climatic Optimum (MECO), originally identified by Bohaty and Zachos (2003) in the Southern Ocean and now recorded globally (Bohaty et al. 2009). The impact of the warming at Hampden was to increase mixedlayer ocean temperature allowing the incursion of tropical foraminifera, deepen or prolong

View facing west of the cliffs towards the southern end of Hampden Beach showing clearly developed

Extended Abstracts


Figure 2 Paleogeographic reconstruction of New Zealand at 40 Ma based on King et al. (1999). The present day coastline is shown in blue.

the seasonal thermocline and increase the ratio of chemical to physical weathering in the terrestrial source region. REFERENCES Bohaty, S.M.; Zachos, J.C. 2003: Significant Southern Ocean warming event in the late middle Eocene. Geology 31: 1017-1020. Bohaty, S.M.; Zachos, J.C.; Floridino, F.; Delaney, M. 2009: Coupled greenhouse warming and deep sea acidification in the Middle Eocene. Paleoceanography 24: PA2207, doi:10.1029/2008PA001676, 2009. Burgess, C.E.; Pearson, P.N.; Lear, C.H.; Morgans, H.E.; Handley, L.; Pancost, R.D.; Schouten, S. 2008: Middle Eocene climate cyclicity in the southern Pacific: Implications for global ice volume. Geology 36: 651-654. Huber, M.; Brinkhuis, H.; Stickley, C.E.; Döös, K.; Sluijs, A.; Warnaar, J.; Schellenberg, S.A.; Williams, G.L. 2004: Eocene circulation of the Extended Abstracts

Southern Ocean: Was Antarctica kept warm by subtropical waters? Paleoceanography 19: PA4026, doi:4010.1029/2004PA001014. King, P.R.; Naish, T.R.; Browne, G.H.; Field, B.D.; Edbrooke, S.W. 1999: Cretaceous to Recent sedimentary patterns in New Zealand. Institute of Geological and Nuclear Sciences folio series 1 Morgans, H.E.G. in press: Late Paleocene–Middle Eocene stratigraphy, foraminiferal biostratigraphy and status of the Bortonian Stage lectostratotype at Moeraki-Hampden coastal section, eastern South Island, New Zealand. New Zealand Journal of Geology and Geophysics. Nelson, C.S.; Cooke, P.J. 2001: History of oceanic front development in the New Zealand sector of the Southern Ocean during the Cenozoic - a synthesis. New Zealand Journal of Geology and Geophysics 44: 535-553.

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Pearson, P.N.; Burgess, C.E. 2008: Foraminifer shell preservation and diagenesis: comparison of high latitude Eocene sites. In Austin, W.A. ed. Biogeochemical Controls on Palaeoceanographic Proxies. Geological Society of London. Shackleton, N.J.; Kennett, J.P. 1975: Paleotemperature history of the Cenozoic and

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the initiation of Antarctic glaciation: oxygen and carbon isotope analyses in DSDP sites 277, 279 and 281. Initial Reports of the Deep Sea Drilling Project 29: 743-755. Zachos, J.C.; Stott, L.D.; Lohmann, K.C. 1994: Evolution of early Cenozoic marine temperatures. Paleoceanography 9: 353-387.

Extended Abstracts


EARLY EOCENE MAMMAL FAUNAL RESPONSE TO TEMPERATURE CHANGE IN THE BIGHORN BASIN, WYOMING Amy Chew Western University of Health Sciences, Pomona CA

INTRODUCTION The Willwood FM in the central part of the Bighorn Basin, WY, has yielded an extraordinarily complete and continuous record of early Eocene mammals (Bown et al. 1994; Rose 2001). More than 1000 fossil localities are known, distributed across ~3000 km2 of fossil-producing badlands. The localities have been tied to a 640 m composite stratigraphic section, which begins at the Paleocene–Eocene (P–E) boundary and ends near the end of the early Eocene. The total time represented by this section is ~3 m.y., during which time two dramatic shifts in mean annual temperature (MAT) occurred following the Paleocene– Eocene Thermal Maximum (PETM) at the P–E boundary. Local MAT dropped suddenly from ~18°C in the 800 k.y. following the PETM to ~10°C and stayed cool for at least 700 k.y. before rebounding back to ~18°C (Wing et al. 2000). The Johns Hopkins/USGS expeditions (1972-present) have collected in excess of 50,000 specimens from the Willwood FM that have been tied to the composite section to form a highly-resolved (no stratigraphic gaps in the fossil sequence >10m) record of faunal evolution. The fossils have been thoroughly studied, making this record an ideal one with which to test mammal faunal response to changing MAT. Previous work (Schankler 1980; Chew 2009) demonstrated significant pulses of fauna-wide change (biohorizons) in taxon richness and rates of turnover that coincided with the shifts in MAT. Biohorizon A was related to the initial decrease in MAT and Biohorizon B was related to the MAT rebound. However, it is unclear whether these events represent shortterm faunal reaction to temperature change or whether community structure was significantly altered in the long-term. Community structure may be modeled by additional paleoecological parameters, including body size and trophic structure. In fossil mammals, relative size is often approximated by the log-transformed occlusal surface area of complete lower first

Extended Abstracts

molars (e.g., Gingerich 1974), while trophic structure is typically represented by the relative frequency and abundance of taxa within trophic groups (e.g., Gunnell et al. 1995). Trophic groups, consisting of taxa with similar life strategies, may be assigned to early Paleogene taxa (for which there are no modern analogues) by an assessment of relative size and dental morphology to roughly determine dietary strategies. Relative size and trophic structure represent important aspects of the organization of early communities that are expected to vary with MAT. For example, the proportions of specialized plant-eating animals should change as the plants they depend on move or disappear with changing MAT. Over long periods of time, body size may also track temperature changes in a manner analogous with Bergmann’s Rule governing modern ecosystems: larger species and/or larger individuals of the same species are expected in cooler conditions. MATERIALS AND METHODS Gnathic and dental specimens were examined by the author in previous work (Chew 2009) and include 41,776 specimens from 621 localities in the Willwood FM, found in the drainage areas of three modern creeks: the Fifteenmile Creek (FMC), the Elk Creek (EC), and the Nowater Creek (NWC) (Table 1). Samples were binned into ~100 k.y. intervals as in Chew (2009). Although fossils from all three regions of the central Bighorn Basin are found in fossil soils (paleosols) and represent autochthonous, attritional lag deposits (Bown and Kraus 1981), there may be preservation differences in the NWC area that have biased the relative size distribution and productivity of individual localities compared with the rest of the basin (e.g., Clyde et al. 2005; Chew and Oheim 2009). A total of 124 lineages were identified and classified following Gunnell et al. (1995) as carnivore, insectivore, omnivore, frugivore, or

21


Figure 1

Proportional richness of the plant-dependent trophic groups.

herbivore. Gunnell et al.’s (1995) herbivore category was further divided into ‘generalists’ and ‘specialists’ for this project on the basis of abundance, relative size and degree of lophodonty (crestiness) of the cheek teeth. Abundant small-medium taxa with low, rounded cheek-tooth cusps were considered generalists, while less abundant, small or large taxa with lophodont cheek teeth were considered to be specialized for consuming particular plant parts. The carnivores, insectivores and omnivores relied on, or supplemented their diets with meat, while the remaining groups were entirely dependent on plants for food. More than half of the total lineages in this sample, 57%, were carnivores, insectivores or omnivores, whereas 88% of the total specimens belonged to the plantdependent groups (68% of the total sample consisted of generalist herbivore specimens). Thirty-three families were identified, all except three (Arctocyonidae, Phenacodontidae and Paromomyidae) of which consisted of lineages with identical trophic strategies. The three exceptional families were subdivided into groups of genera with consistent dietary strategies. As sample sizes of the 100 k.y. intervals varied by more than two orders of magnitude (Table 1), extensive standardization of lineage richness and abundance was necessary for the assessment of trophic structure. Lineage richness was rarefied to a standardized sample size of 50 specimens. The ‘average proportional diversity’ of each family (or group of genera in the case of the exceptional

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three families) was calculated from the bootstrapped proportion of lineages compiled by family/generic group out of the total lineage richness per 100 k.y. interval following Alroy (2000). This statistic combines both the abundance and richness aspects of diversity into a single measure reflecting the relative importance of the individual families/generic groups at any given time. As each family/generic group consists of lineages from a single trophic category, this statistic summarizes trophic structure. Relative size was estimated by log-transformed occlusal surface areas of all complete lower first molars, averaged by 100 k.y. interval within lineages that spanned the entire section. Both proportional diversity and relative sizes were contrasted by cool versus warm periods. RESULTS AND DISCUSSION Trophic structure. As expected, the proportional diversity curves of the plantdependent trophic groups were the only ones to show unequivocal changes in concert with MAT (Fig. 1). The specialized herbivorous families, including the ancestors of modern rodents (Paramyidae) and tapirs (Isectolophidae), as well as some archaic groups with lophodont teeth, reached their peak proportionate diversities during the cool period. In total, specialized herbivores made up an average of 30% of the lineages present in samples during the cool period, which is more than any of the other combined trophic groups. In contrast, generalized herbivorous and frugivorous families had relatively static or

Extended Abstracts


Interval (Ma)

Meter levels

Area

MAT

N

55.79-55.7

0-23

NWC

Warm

38

55.69-55.6

24-46

NWC

Warm

1338

55.59-55.5

47-69

NWC

Warm

623

55.49-55.4

70-92

NWC

Warm

212

55.39-55.3

93-115

NWC+EC

Warm

340

55.29-55.2

116-137

NWC+EC

Warm

64

55.19-55.1

138-160

EC

Warm

981

55.09-55.0

161-183

EC

Warm

201

54.99-54.9

184-206

EC

Biohorizon A Cooling?

360

54.89-54.8

207-229

EC

Cool

243

54.79-54.7

230-251

EC

Cool

483

54.69-54.6

252-274

EC+FMC

Cool

553

54.59-54.5

275-297

EC+FMC

Cool

574

54.49-54.4

298-320

EC+FMC

Cool

386

54.39-54.3

321-343

EC+FMC

Cool

1010

54.29-54.2

344-365

EC+FMC

Cool

1547

54.19-54.1

366-388

EC+FMC

495

54.09-54.0

389-411

EC+FMC

Biohorizon B Warming?

3190

53.99-53.9

412-434

EC+FMC

Warm

2432

53.89-53.8

435-455

FMC

Warm

3870

53.79-53.7

456-470

FMC

Warm

4153

53.69-53.6

471-485

FMC

Warm

4889

53.59-53.5

486-500

FMC

Warm

3851

53.49-53.4

501-515

FMC

Warm

1208

53.39-53.3

516-530

FMC

Warm

864

53.29-53.2

531-545

FMC

Warm

708

53.19-53.1

546-560

FMC

Warm

6078

53.09-53.0

561-575

FMC

Warm

221

52.99-52.9

576-589

FMC

Warm

43

52.89-52.8

590-604

FMC

Warm

736

52.79-52.7

605-619

FMC

Warm

0

52.69-52.6

620-634

FMC

Warm

85

Table 1 Temperature and sampling distribution of the 100Kyr intervals. N: Total number of specimens per 100Kyr interval. MAT: Mean annual temperature. Area: Interval includes localities from the Nowater Creek (NWC), Elk Creek (EC) and/or Fifteenmile Creek (FMC) drainage basins.

declining proportional diversities during the cool period, with a combined proportional diversity never exceeding 27% of the lineages present. Only one generalist herbivore family, the Dichobunidae, experienced a slight increase in proportional diversity during the cool interval because of the addition of a

Extended Abstracts

larger, more robust lineage during this time. In the following warm period, the proportional diversities of the generalist herbivores and frugivores suddenly began to increase, reaching their peaks towards the end of the section when MAT was highest. By the end of the section, the two groups combined made up 61.5% of the lineages present in the sample. The generalist herbivores and frugivores were not particularly lineage-rich during the initial warm period. However, much of this period (55.8–55.2Ma) is represented by specimens from the NWC area where unique conditions have probably led to the preservation of a greater proportion of small, relatively rare taxa (Chew 2009). Although absolute abundances are difficult to assess for fossil assemblages with sample sizes -10‰) are typical for aquatic consumers foraging in estuaries and marine foodwebs (phytoplankton-, kelp-, or seagrass-based) or terrestrial consumers foraging in C4 grasslands. These separations can be further refined by identification of aquatic consumers through oxygen isotope analysis. Analysis of living semi-aquatic and aquatic Figure 1 Phylogeny of the Cetacea based on Thewissen mammals, particularly hippopotamids, has et al. (2007) and Uhen (2008). Specimens of families listed in bold were included in this study.

26

Extended Abstracts


shown that these species often yield lower enamel δ18O values than those for the associated terrestrial fauna (Bocherens et al. 1996; Clementz et al. 2008; Levin et al. 2006). This difference is a result of a greater influx and efflux of surface water passing through the bodies of aquatic species than for terrestrial species (Clementz et al. 2008). Using this relationship along with enamel δ13C values, it is possible to determine habitat of an extinct species, and whether it foraged in freshwater or marine foodwebs. Prior research has applied these relationships between geochemistry and ecology to answer questions about the diets and habitat of early whales (Clementz et al. 2006; Roe et al. 1998), but these studies were restricted to Cetacea and did not include sister taxa. In addition, the amount of archaeocete material available for stable isotope analysis was limited. Discoveries of multiple well-preserved specimens of archaeocetes and artiodactyls from India and Pakistan now makes it feasible to expand isotope work and look at how the feeding ecology of cetaceans shifted from the artiodactyl norm during the earliest stages of their evolution to that of aquatic predators. METHODS Teeth were collected and sampled from two species of raoellid, the sister group to Cetacea, and eight archaeocete species (listed in bold), which were then compared to published data for eight species from four of the five archaeocete families (Table 1; Fig. 1). Archaeocete material was collected from several localities in India, Pakistan, Egypt and North America (Fig. 2), spanning a considerable range of ages (Early to Late Eocene) and depositional environments (freshwater to marine). Approximately 10 mg of enamel powder were collected from the base of each sampled tooth and chemically prepared for stable isotope analysis following standard procedures (Koch et al., 1997). All isotope values are reported in standard delta notation, where δ= ((Rsample/Rstandard) – 1) x 1000 and R is 13 12 C/ C for carbon and 18O/16O for oxygen. Carbon values are reported relative to the VPDB standard and δ18O values are reported relative to standard mean ocean water (VSMOW). Precision of isotopic analysis performed at the University of Wyoming Extended Abstracts

Stable Isotope Facility (UWSIF) was assessed via multiple analyses of in-house elephant enamel and fossilized sirenian bone standards (δ13C: standard deviation s = 0.1‰; δ18O: s = 0.2‰; n = 30 for both). RESULTS AND DISCUSSION Enamel stable isotope values indicate substantial ecological, environmental and, particularly for early whales, evolutionary change through the Eocene. For the raoellid Indohyus indirae, enamel δ18O values were extremely low and significantly depleted in δ18O relative to coeval terrestrial mammals from the same locations (Table 1). These low values are indicative of semiaquatic, freshwater habits and suggest an early movement into the water by the common ancestor of this group and whales, possibly at the very start of the Eocene. Carbon isotope values for Indohyus, however, do not support an aquatic diet for these early whale relatives, which may mean that food was not the initial motivation for whales to take to the water. These findings were reported in the December 2007 issue of Nature. Subsequent analysis of enamel from related artiodactyls (Raoellidae: Khirtharia) has shown similar results; extremely low δ18O values distinct from those of other mammals in the fauna, but enamel δ13C values that suggest a land-based diet. As more material from artiodactyls (dichobunids) and other herbivores (cambaytheriids) becomes available, enamel stable isotope values for these species will provide a more refined interpretation of the diet variation within Early Eocene ungulates and, more specifically, the sister family to cetaceans. With the appearance of the earliest cetaceans, however, there is strong stable isotope evidence that these species were fully exploiting freshwater foodwebs. As with the raoellids, enamel δ18O values for these early whales are extremely low and suggest a semiaquatic lifestyle with considerable time spent in freshwater ecosystems. Unlike raoellids, however, enamel carbon isotope values for the pakicetids Pakicetus inachus, Pakicetus attocki, Nalacetus ratimitus and Icthyolestes pinfoldi are extremely low for these species and consistent with a fish-based diet fuelled by freshwater phytoplankton (Table 1; Fig. 3). Values for these species are 27


significantly higher than reported mean values for most living freshwater cetaceans, but are similar to values for river otters foraging in rivers and streams (Clementz, unpublished data). High carbon isotope values in freshwater foodwebs are typically a sign of well-mixed waters and may imply that pakicetids were foraging primarily in streams and rivers rather than stagnant pools, ponds or lakes.

Table 1

28

Analysis of tooth enamel from the more derived and aquatically mobile Ambulocetidae (Ambulocetus natans) shows at least one species in this family had a similar affinity for freshwater habitats. These findings are surprising and in direct opposition to the sedimentary evidence associated with these fossils (Williams 1998). Ambulocetid remains are typically found in deltaic or shallow marine sediments, which suggest a preference for

Mean δ13C and δ13C values for specimens of Raoellidae and Archaeoceti. Extended Abstracts


Figure 2 Early Eocene paleogeographic reconstruction showing locations (yellow circles) from which fossil cetacean and raoellid specimens sampled for this study were taken. Map downloaded from the Paleobiology Database, March 2009.

coastal habitats. However, enamel δ13C and δ18O values are extremely low and nearly identical to values for freshwater pakicetids, which support a freshwater diet and habitat for this family. Analysis of specimens from other species of this family (Gandakasia) is planned to see if all ambulocetids were restricted to freshwater habitats. These findings imply a prolonged freshwater stage in cetacean evolution (ca 53.5–48.0 Ma) and suggest that the increased body size of ambulocetids relative to pakicetids was not facilitated by the transition to marine ecosystems. Subsequent sampling of later whales shows evidence of marine foraging by the early middle Eocene (ca 48 Ma), but with a preference for shallow, nearshore habitats. Specimens from the families Remingtonocetidae and Protocetidae exhibit increasing aquatic specialization and fossil remains for these families are only found in marine sediments. Enamel carbon and oxygen isotope values for these families are consistently elevated with respect to those for pakicetids and ambulocetids and fall close to expected marine consumer values based on data from living marine species (Fig. 3).

Extended Abstracts

Extremely high δ13C values with little individual variation for Attockicetus, Remingtonocetus and Andrewsiphius indicate foraging in nearshore, coastal waters and show that this preference developed early among members of this family.

Figure 3 Bivariate plot of carbon and oxygen isotope values for fossil cetacean and raoellid tooth enamel. For taxa of sufficient specimen number (n ≥3), mean values are reported and error bars represent ±1 SD. Specimens are grouped by family and identified by colour (Raoellidae: yellow; Pakicetidae: light green; Ambulocetidae: dark green; Remingtonocetidae: light blue; Protocetidae: blue; Basilosauridae: dark blue; living Odontoceti: white). Gray diamonds represent values for fossil bones and teeth reported in Roe et al. (1998).

29


Species of basilosaurids also have high δ13C and δ18O enamel values typical for marine consumers. Based on the strong negative correlation between latitude and the oxygen and carbon isotope compositions of seawater today (Bigg et al. 2000; Rau et al. 1989; Schmidt 1999; Williams 1998), the high stable isotope values for Basilosaurus and Dorudon are consistent with low latitude foraging by these species and imply that they did not migrate significant distances to higher latitude through the year. However, similar values for protocetids and basilosaurids sampled from localities in India, Africa and North America do not preclude the possibility that these species may have travelled significant longitudinal distances while remaining at low latitudes when feeding.

Cerling, T. E.; Harris, J. M.; Macfadden, B. J.; Leakey, M. G.; Quade, J.; Eisenmann, V.Ehleringer, J. R. 1997: Global vegetation change through the Miocene/Pliocene boundary. Nature 389: 153-158

CONCLUSIONS

Clementz, M. T.; Goswami, A.; Gingerich, P. D.; Koch, P. L. 2006: Isotopic records from early whales and sea cows: Contrasting patterns of ecological transition. Journal of Vertebrate Paleontology 26: 355-370

Stable isotope analysis of tooth enamel from the Archaeoceti and the Raoellidae — fossil sister group to Cetacea — document a significant ecological shift in diet and habitat preferences that parallels morphological change associated with the transition from terrestrial to aquatic ecosystems. Based on coupled carbon and oxygen isotope analyses, the transition to an aquatic existence appears to have preceded a switch to an aquatic diet, though this change in foraging preferences occurred only a short time later. Once in the water, early whales initially foraged in freshwater food webs before moving out into marine waters and refining morphological adaptations for swimming. Stable isotope data for the most derived archaeocetes indicates these species foraged primarily in productive, low latitude waters near to shore, which implies that foraging offshore or at high latitudes did not occur until later with the Neoceti (odontocetes, mysticetes). REFERENCES Bigg, G. R.; Rohling, E. J. 2000: An oxygen isotope data set for marine water. Journal of Geophysical Research 105: 8527-8535 Bocherens, H.; Koch, P.; Mariotti, A.; Geraads, D.; Jaegar, J.-J. 1996: Isotopic biogeochemistry (δ13C, δ18O) of mammalian enamel from African Pleistocene hominid sites. Palaios 11: 306-318

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Cerling, T. E.; Harris, J. M. 1998: Carbon isotopes in bioapatite of ungulate mammals: Implications for ecological and paleoecological studies. Journal of Vertebrate Paleontology 18: 33A. Clementz, M. T.; Koch, P. L. 2001: Differentiating aquatic mammal habitat and foraging ecology with stable isotopes in tooth enamel. Oecologia 129: 461-472 Clementz, M. T.; Hoppe, K. A.; Koch, P. L. 2003: A paleoecological paradox: the habitat and dietary preferences of the extinct tethythere Desmostylus, inferred from stable isotope analysis. Paleobiology 29: 506-519

Clementz, M. T.; Holroyd, P. A.; Koch, P. L. 2008: Identifying aquatic habits of herbivorous mammals through stable isotope analysis. Palaios 23: 574-585 Geisler, J. H.; Uhen, M. 2005: Phylogenetic relationshipbs of extinct cetartiodactyls: results of simultaneous analyses of molecular, morphological, and stratigraphic data. Journal of Mammal Evolution 12: 145-160 Hoppe, K. A.; Koch, P. L.; Carlson, R. W.; Webb, S. D. 1999: Tracking mammoths and mastodons: reconstruction of migratory behavior using strontium isotope ratios. Geology 27: 439-442 Koch, P. L.; Fisher, D. C.; Dettman, D. 1989: Oxygen isotope variation in the tusks of extinct proboscideans: a measure of season of death and seasonality. Geology 17: 515-519 Levin, N. E.; Cerling, T. E.; Passey, B. H.; Harris, J. M.; Ehleringer, J. R. 2006: A stable isotope aridity index for terrestrial environments. Proceedings of the National Academy of Science 103: 11201-11205 Macfadden, B. J.; Higgins, P.; Clementz, M. T.; Jones, D. S. 2004: Diets, habitat preferences, and niche differentiation of Cenozoic sirenians from Florida: evidence from stable isotopes. Paleobiology 30: 297-324 O'leary, M. A.; Uhen, M. D. 1999: The time of origin of whales and the role of behavioural

Extended Abstracts


changes in the terrestrial-aquatic transition. Paleobiology 25: 534-556 Rau, G. H.; Takahashi, T.; Des Marals, D. J. 1989: Latitudinal variations in plankton d13C: implications for CO2 and productivity in past oceans. Nature 341: 516-518 Schmidt, G. A. 1999: Forward modeling of carbonate proxy data from planktonic foraminifera using oxygen isotope tracers in a global ocean model. Paleoceanography 14: 482-497

Extended Abstracts

Thewissen, J. G.; Bajpai, S. 2001: Whale origins as a poster child for macroevolution. BioScience 51: 1017-1029 Zachos, J. C.; Pagani, M.; Sloan, L. C.; Thomas, E.; Billups, K. 2001: Trends, rhythms, and aberrations in global climate 65 Ma to present. Science 292: 686-693 Zachos, J. C.; Dickens, G. R.; Zeebe, R. E. 2008: An early Cenozoic perspective on greenhouse warming and carbon-cycle dynamics. Nature 451: 279-83

31


PERTURBATIONS TO THE GLOBAL C-CYCLE AT THE PETM AND DURING THE TOARCIAN (EARLY JURASSIC) Anthony S. Cohen and Angela L. Coe Department of Earth and Environmental Sciences, The Open University, MK7 6AA, UK: [email protected]

INTRODUCTION Environmental change is a natural feature of Earth’s dynamic state; change is caused by many different processes that operate over a wide range of temporal and spatial scales. The relatively rapid fluctuations in global climate that have occurred over the last ~3 Ma of Earth history are generally attributed to variations in insolation, resulting in the many glacialinterglacial cycles of the Quaternary. However, on relatively rare occasions in Earth’s more distant past, transient ‘hyperthermal’ episodes are thought to have been caused by sudden increases in the level of atmospheric CO2 (Jenkyns 2003; Zachos et al. 2001). The bestknown example of a hyperthermal resulting from a sudden increase in atmospheric CO2 in the Paleogene was the relatively brief episode of extreme warmth at the Paleocene–Eocene boundary, which is usually referred to as the Paleocene–Eocene Thermal Maximum (PETM) (Zachos et al. 2001). In the Mesozoic, transient hyperthermals that were likely to have resulted from similar processes occurred on at least three separate occasions. These Mesozoic hyperthermals were also associated with so-called Oceanic Anoxic Events (OAEs) (Jenkyns 2003). At the present-day, anthropogenic global warming is taking place mainly because of the rapid introduction to the atmosphere of unusually large quantities of CO2. Whilst this scenario is in many respects notably similar to those of the (relatively rare) hyperthermals that were driven by large-scale CO2 release, it differs greatly from the insolation-controlled glacial-interglacial cycles of the Quaternary. The records of these ancient, CO2-driven hyperthermals may thus provide us with very useful analogues for anthropogenic global warming. DISCUSSION The PETM and the 3 hyperthermals of the Mesozoic, which were associated with major OAEs (the Early Toarcian OAE, the ‘Selli’ vent (OAE1a) and the ‘Bonarelli’ event

32

(OAE2)), were also associated with major perturbations to the global carbon cycle. For example, the records of these events provide evidence of seawater acidification, pronounced carbon isotope excursions (CIEs) in marine and terrestrial carbon reservoirs, enhanced accumulation of organic-rich sediments, increased levels of seawater dysoxia and anoxia, and mass extinctions of marine and terrestrial species (Erba 2004; Hesselbo et al. 2000; Pearce et al. 2008; Thomas 2007; Zachos et al. 2001). Of particular note is the observation that the pattern and duration of the Toarcian CIE, as preserved in both marine organic carbon and carbonate, is remarkably similar to the key features of the CIE that characterises the PETM (Cohen et al. 2007). The similarities between the two events are important because they suggest that the Earth System responded in a consistent manner despite the fact that the events were separated by ~124 Ma and that they occurred under very different background conditions of global climate and Paleocontinental configuration. In other words, although both events were rare, neither was unique. A corollary of this finding is that it should in principle be possible to use observations from one event to understand aspects of the other that may be less well defined. Accurate timescales for both the PETM and Toarcian CIEs allow us to estimate the rates at which carbon was added to the oceanatmosphere system during the hyperthermals and the speed at which the events were established. This information also enables us to define the overall durations of the events. For the PETM, a range of estimates based upon different techniques yield relatively consistent results. The identification of astronomical cycles (Norris and Röhl 1999; Röhl et al. 2007) and measurements of 3He accumulation rates (Farley and Eltgroth 2003) indicate that the event lasted between ~170 and ~220 k.y. following the first abrupt carbon isotope shift. The astronomical timescale of Kemp et al. (2005) was based upon the analysis of samples

Extended Abstracts


from the Early Jurassic sections of the Yorkshire coast in the U.K., and indicates that the main phase of the CIE lasted ~225 k.y. after the first abrupt carbon isotope shift (Cohen et al. 2007). A somewhat longer duration of ~600 k.y. was suggested by Suan et al. (2008), based on the analysis of samples from a section in Portugal. However, this section contains turbidites and its precise stratigraphic correlation with the U.K. sections is unclear. An important feature that characterises the CIEs at both the PETM and in the Early Toarcian is that they are built up from multiple, sudden and substantial shifts to negative δ13C values that affected the entire biospheric carbon reservoirs both on land and in the oceans (see Cohen et al. (2007) for a review). The presence of a pronounced carbonate dissolution horizon that occurs at the onset of the CIE in many PETM sections (e.g. Zachos et al. 2005) often makes it hard to obtain δ13C(carb) data that indicate reliably the magnitude and duration of the initial δ13C shift. However, carbon isotope data obtained from organic carbon (both bulk organic matter and specific compounds) are not affected by dissolution and they thus have the potential to provide reliable records of the CIE. The data reported by Kemp et al. (2005) for the Toarcian CIE exemplify this point and demonstrate the rapidity of the initial shifts in δ13C. Based upon the precessional timescale derived by Kemp et al. (2005), the first two δ13C shifts of ~3‰ occurred in no more than ~2 k.y. and the second two shifts of ~2‰ took place in as little as ~650 years (Cohen et al. 2007). A recent study based on the analysis of relatively shallow marine Toarcian carbonates from the Paris basin in France reported a similar pattern for the CIE, again comprising four abrupt shifts each of ~2‰ to ~3‰ (Hermoso et al. 2009). It is noteworthy that for the PETM, the magnitude of the CIE in samples of both marine organic carbon and terrestrial pedogenic carbon is often larger than that seen in marine carbonate (Bowen et al. 2004; Pagani et al. 2006). Whilst the source of the isotopically ‘light’ carbon that produced these abrupt shifts has been the subject of much debate, a diverse range of observations constrains rather tightly the mechanisms that could have produced Extended Abstracts

them. In particular, the magnitude and rapidity of the δ13C shifts together with their evident astronomical pacing (Kemp et al. 2005; Lourens et al. 2005) require a large and readily available supply of ‘light’ carbon that could have become activated by astronomicallycontrolled fluctuations in climate and/or marine conditions. The sudden dissociation of large masses of methane hydrate, as suggested originally by Dickens et al. (1995), provides a mechanism that readily fits the observations and constraints for the main phases of the CIEs during both the PETM and the Toarcian. Mass balance calculations suggest that ~2,000 Gt methane would have been required to produce each δ13C shift of ~2–3‰ (Dickens et al. 1995; Beerling and Brentnall 2007). Under the assumption that the carbon source was methane hydrate, the two upper δ13C shifts of ~2‰ that each occurred over a maximum of ~650 years during the Toarcian CIE would have each required an average release rate of ~4 Gt carbon a-1. This figure is a minimum and is ~50% of the present-day release rate of anthropogenic carbon to the atmosphere. The expansion of marine anoxia during the Mesozoic OAEs has, until recently, been inferred entirely from field observations which suggest that the accumulation of organic carbon was enhanced significantly during those intervals (e.g. Schlanger and Jenkyns 1976). Recently, Pearce and colleagues (Pearce et al. 2008) reported the first Mo-isotope data for any OAE, based upon Toarcian samples from the U.K. Where the Mo is derived predominantly from euxinic seawater, this new isotopic proxy is able to reflect changes in the areal extent of marine anoxia worldwide. The Mo-isotope data reported by Pearce et al. (2008) indicate that the areal extent of marine anoxia fluctuated by at least an order of magnitude on a precessional timescale throughout the Toarcian CIE, broadly in line with the previously reported changes in δ13C (Kemp et al. 2005). At the maximum extent of anoxia during the Toarcian CIE, an area equivalent to most of the world’s continental shelf appears to have become anoxic. During the PETM, there was widespread accumulation of organic-rich strata in many shallow seas in the northern hemisphere (Gavrilov et al. 1997) and conditions in many other continental shelf areas may have become at least dysoxic. The very obvious similarities between the PETM 33


and the Toarcian CIEs prompted Cohen et al. (2007) to suggest that the PETM may have been an incipient OAE. Future geochemical studies should help to define the extent to which marine anoxia did in fact change during the PETM. REFERENCES Beerling, D.J.; Brentnall, S.J. 2007: Numerical evaluation of mechanisms driving Early Jurassic changes in global carbon cycling. Geology 35: 247-250. Bowen, G.J.; Beerling, D.J.; Koch, P.L.; Zachos, J.C.; Quattlebaum, T. 2004: A humid climate state during the Paleocene/Eocene thermal maximum. Nature 432: 495-499. Cohen, A.S.; Coe, A.L.; Kemp, D.B. 2007: The Late Paleocene Early Eocene and Toarcian (Early Jurassic) carbon isotope excursions: a comparison of their time scales, associated environmental changes, causes and consequences. Journal of the Geological Society 164: 1093-1108. Dickens, G. R.; Oneil, J.R.; Rea, D.K.; Owen, R.M. 1995: Dissociation of Oceanic Methane Hydrate as a Cause of the Carbon-Isotope Excursion at the End of the Paleocene. Paleoceanography 10: 965-971. Erba, E. 2004: Calcareous nannofossils and Mesozoic oceanic anoxic events. Marine Micropaleontology 52: 85-106. Farley, K.A.; Eltgroth, S.F. 2003: An alternative age model for the Paleocene-Eocene thermal maximum using extraterrestrial 3He. Earth Planetary Science Letters 208: 135-148. Gavrilov, Y.O.; Kodina, L.A.; Lubchenko, I.Y.; Muzylev, N.G. 1997: The Late Paleocene Anoxic Event in Epicintinental Seas of PeriTethys and Formation of the Sapropelite Unit: Sedimentology and Geochemistry. Lithology and Mineral Resources 32: 427-450. Hesselbo, S. P.; Gröcke, D. R.; Jenkyns, H.C.; Bjerrum, C.J.; Farrimond, P.; Bell, H.S.M.; Green, O.R. 2000: Massive dissociation of gas hydrate during a Jurassic oceanic anoxic event. Nature 406: 392-395. Jenkyns, H.C. 2003: Evidence for rapid climate change in the Mesozoic-Palaeogene greenhouse world. Philosophical Transactions of the Royal Society 361:1885-1916.

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Kemp, D.B.; Coe, A.L.; Cohen, A.S.; Schwark, L. 2005: Astronomical pacing of methane release in the Early Jurassic period. Nature 437: 396399. Lourens, L.J.; Sluijs, A.; Kroon, D.; Zachos, J.C.; Thomas, E.; Rohl, U.; Bowles, J.; Raffi, I. 2005: Astronomical pacing of late Palaeocene to early Eocene global warming events. Nature 435: 1083-1087. Norris, R.D.; Rohl, U. 1999: Carbon cycling and chronology of climate warming during the Palaeocene/Eocene transition. Nature 401: 775-778. Pagani, M.; Pedentchouk, N.; Huber, M.; Sluijs, A.; Schouten, S.; Brinkhuis, H.; SinningheDamste, J.; and the IODP Expedition 302 Scientists 2006: Arctic hydrology during global warming at the Palaeocene/Eocene thermal maximum. Nature 442: 671-675. Pearce, C. R.; Cohen, A. S.; Coe, A. L.: Burton, K. W. 2008: Molybdenum isotope evidence for global ocean anoxia coupled with perturbations to the carbon cycle during the Early Jurassic. Geology 36: 231-234. Röhl, U.; Westerhold, T.; Bralower, T.J.; Zachos, J.C. 2007: On the duration of the PaleoceneEocene thermal maximum (PETM). Geochemistry, Geophysics, and Geosystems 8: doi: 10.1029/2007GC001784. Schlanger, S.O.; Jenkyns, H.C. 1976: Cretaceous Oceanic Anoxic Events: Causes and Consequences. Geologie en Mijnbouw 55: 179184. Thomas, E. 2007: Cenozoic mass extinctions in the deep sea; what disturbs the largest habitat on Earth? In S. Monechi, R. Coccioni and M. Rampino ed. Large Ecosystem Perturbations: Causes and Consequences. Geological Society of America. Pp. 1-24. Zachos, J.; Pagani, M.; Sloan, L.; Thomas, E.; Billups, K. 2001: Trends, rhythms, and aberrations in global climate 65 Ma to present. Science 292: 686-693. Zachos, J. C.: Röhl, U.; Schellenberg, S.; Sluijs, A.; Hodell, D.; Kelly, D.C.; Thomas, E.; Nicolo, M.; Raffi, I.; Lourens, L.; McCarren, H.; Kroon, D. 2005: Rapid acidification of the ocean during the Paleocene-Eocene thermal maximum. Science 308: 1611-1615.

Extended Abstracts


OVERVIEW OF MULTIPLE ENVIRONMENTAL SIGNALS ACROSS THE ONSET OF THE PALEOCENE–EOCENE THERMAL MAXIMUM (PETM) IN THE COBHAM LIGNITE BED (KENT, ENGLAND) M.E.Collinson1, D. Steart1,2,3, R.D. Pancost4, J.J. Hooker2, G.J. Harrington5, A.C.Scott1, L. Handley4, I.J. Glasspool6, L.O. Allen7 and S.J. Gibbons1 1

Department of Earth Sciences, Royal Holloway University of London, Egham, Surrey TW20 0EX, UK: [email protected]; 2Department of Palaeontology, The Natural History Museum, Cromwell Road, London, SW7 5BD, UK; 3Department of Palaeontology, University of Witwatersrand 2050, South Africa; 4 Organic Geochemistry Unit, Bristol Biogeochemistry Research Centre, School of Chemistry, University of Bristol, Cantocks Close, Bristol, BS8 1TS, UK; 5Department of Geography, Earth and Environmental Sciences, University of Birmingham, Birmingham, B15 2TT, U.K.; 6Department of Geology, The Field Museum, 1400 S lake Shore Drive, Chicago, Illinois 60605-2496, USA; 7Burnet Close, Hemel Hempstead, Hertfordshire, UK.

The Cobham Lignite Bed occurs in a small outlier of Cenozoic deposits in the London Basin, located at Scalers Hill, near Cobham, Kent, England (Fig. 1 in Steart et al. 2007). The stratigraphy, relative dating and regional correlation were given in Collinson et al. (2003) and updated calibration of PETM onset was provided in Collinson et al. (2009). The Cobham Lignite Bed is underlain by a sand and mud unit, which is itself underlain, disconformably, by the Upnor Formation. The Cobham Lignite Bed is overlain, successively, by lower shelly clays of the Woolwich Formation, Blackheath Beds, Oldhaven Member and London Clay Formation. The Cobham Lignite Bed itself comprises a lower laminated lignite and an upper blocky lignite. Thin clays occur at the base and between the two lignites. A unique suite of paleoenvironmental signals across the onset of the PETM has been revealed by qualitative and quantitative coal petrological, plant mesofossil, palynological and compound-specific carbon-isotope analyses of the Cobham Lignite Bed. A negative shift (CIE) in δ13C values of the bulk lignite and C29 and C27 n-alkanes confirms the presence and stratigraphic location of the PETM onset near the top of the laminated portion of the Cobham Lignite Bed (Collinson et al. 2003; Pancost et al. 2007). Based on peat to coal compaction ratios the overlying blocky lignite represents an estimated 4–12 k.y. after PETM onset. The PETM continued through the subsequent deposition of lower shelly clays of the Woolwich Formation (lower Woolwich Shells Beds), which contain the Apectodinium acme.

Extended Abstracts

At PETM onset there is a striking and unique negative CIE in bacterially-derived hopanes (Pancost et al. 2007). This indicates an increase in the methanotroph population and suggests positive feedback in response to a warming climate with increased precipitation, leading to increased release of methane from the terrestrial biosphere (Pancost et al. 2007). An increasing level of precipitation across the PETM is supported by major changes in both the vegetation and fire regime that are coincident with this interval (Collinson et al. 2007, 2009; Steart et al. 2007). The palynomorph composition of Late Paleocene samples is significantly different from PETM samples (Collinson et al. 2009). Prior to PETM onset the Cobham laminated lignite has charcoal-rich and charcoal-poor layers, indicative of episodic fires and post fire erosion. Herbaceous fern leaf stalks and angiosperm wood fragments dominate the charcoal clasts (Collinson et al. 2007; Steart et al. 2007). Cicatricosisporites (Schizaeaceae) fern spores, microscopic and mesoscopic charcoal co-occur in high abundances (Collinson et al. 2007, 2009). The source vegetation was a low diversity, fire prone, fern and woody angiosperm community, adapted to fire disturbance and seasonal surface wildfires. In contrast, post-PETM onset, the blocky lignite is derived from decomposed plant material (peat) and lacks charcoal showing that fires have ceased (Steart et al. 2007). There is a more varied flowering plant community with palms and eudicots and an increase in wetland plants especially taxodiaceous conifers (Collinson et al. 2009). Mesofossils from a free-floating aquatic plant community also occur in the blocky lignite. These include the heterosporous water ferns Azolla and Salvinia 35


represented by megaspores and microspore massulae, described originally by Martin (1976) but now localised to within the PETM. Raw and rarefied palynomorph species richness measures are higher in the PETM than pre-PETM but the difference is not statistically significant (Collinson et al. 2009). Only five (of 24) common palynomorph taxa have last appearance or major shifts in percentage occurrence close to the PETM onset (Collinson et al. 2009). One of these, a triporate eudicot (Triporopollenites 1 of Collinson et al. 2009), occurs only in two samples (SH24 and 25), with extremely high absolute numbers and high percentages in sample 25 (Collinson et al. 2009). Sample 25 is the oldest in the group of three samples (25, 26, 27) with δ13C bulk lignite values of -27 to -27.15‰ at the base of the CIE. δ13C hopane values already exhibit a negative shift in sample 24 (-45‰) becoming more negative in sample 25 (-56‰) and 26 (65‰) and reaching their maximum negative δ13C value in sample 27 (-76‰). Samples 24 to 27 span an interval of only 11.5 mm of sediment thickness in the upper part of the laminated lignite. The occurrence of Triporopollenites 1 is therefore an interesting, but unexplained, feature of PETM onset. The pollen is not only unusual for its stratigraphic occurrence and abundance but also because it is dark brown to black, a very different colour from all other pollen in the samples, (yellow to orange-brown) (Collinson et al. 2009). The taxonomic affinity is not known. The overall grain morphology and the distribution of microechinae seen under the SEM (not identical to that of modern families such as Juglandaceae, Betulaceae or Myricaceae, e.g. Moore et al. 1991) suggest that it may have been produced by an extinct taxon in the Fagales. This pollen may have potential as a marker for PETM onset if it can be identified in other European successions. ACKNOWLEDGEMENTS We thank the Leverhulme Trust for providing funding (Grant number F/07/537/0); Alfred McAlpine, AMEC and Channel Tunnel Rail Link for access to the Cobham site; Sharon Rose for making arrangements on site and Jackie Skipper and Steve Tracey for help with initial sample collection and field discussions.

36

We thank Steve Tracey for providing most samples of the lower shelly clays of the Woolwich Formation from which the occurrence of the Apectodinium acme was determined. We thank Alex Wright, who undertook a study of the Apectodinium in part fulfilment of the degree of MSci in Royal Holloway University of London, and Appy Sluijs for confirming acme percentages of Apectodinium in the Woolwich Formation at Cobham. REFERENCES Collinson, M.E.; Hooker, J.J.; Grocke, D.R. 2003: The Cobham Lignite Bed and penecontemporaneous macrofloras of S. England: a record of vegetation and fire across the Paleocene-Eocene thermal maximum. Geological Society of America Special Paper 369: 333-349. Collinson, M.E.; Steart, D.; Scott, A.C.; Glasspool, I.J.; Hooker, J.J. 2007: Episodic fire, runoff and deposition at the Paleocene-Eocene boundary. Journal of the Geological Society of London 164: 87-97. Collinson, M.E.; Steart, D.; Harrington, G.J.; Hooker, J.J.; Scott, A.C.; Allen, L.O.; Glasspool, I. J.; Gibbons, S.J. 2009: Palynological evidence of vegetation dynamics in response to palaeoenvironmental change across the onset of the Paleocene-Eocene Thermal maximum at Cobham, Southern England. Grana 48: 38-66. Martin, A.R.H. 1976: Upper Paleocene Salviniaceae from the Woolwich/Reading Beds near Cobham, Kent. Palaeontology 19: 173184. Moore, P.D.; Webb, J.A.; Collinson, M.E. 1991: Pollen Analysis (second edition). Oxford, Blackwell Scientific Publications. 216 p. Pancost, R.D.; Steart, D.S.; Handley, L.; Collinson, M.E.; Hooker, J.J.; Scott, A.C.; Grassineau, N.V.; Glasspool, I.J. 2007: Increased terrestrial methane cycling at the Palaeocene-Eocene thermal maximum. Nature 449: 332-335. Steart, D.C.; Collinson, M.E.; Scott, A.C.; Glasspool, I.J.; Hooker, J. 2007: The Cobham Lignite Bed: the palaeobotany of two petrographically contrasting lignites from either side of the Paleocene-Eocene carbon isotope excursion. Acta Palaeobotanica 47: 109-125.

Extended Abstracts


DEPOSITIONAL HISTORY OF THE PALEOGENE ADRIATIC CARBONATE PLATFORM V. Ćosović1, K. Drobne2 and A. Moro1 1

Department of Geology, Faculty of Science, Horvatovac 102a, Zagreb, Croatia: [email protected]; 2I. Rakovec Institute of Paleontology, ZRC SAZU, Novi trg 2, Ljubljana, Slovenia.

The Paleogene Adriatic carbonate platform (PgAdCP) existed within the Central Tethys (around 32°N Paleolatitude) from the Paleocene (Danian) to late Middle Eocene (Bartonian). During this time, the PgAdCP was elongated in a NW–SE trending gulf opening to the south, west and east. Its depositional history differs from the pattern of stages recognized in the Paleocene to the Early Eocene platform development in the Pyrenees and Egypt (Scheibner and Speijer 2008). The differences are the result of delayed generation of a shallow-marine regime (due to tectonic movements such as the uplifts of the Dinarides, Apennines and Alps) and development of different bioconstructers within certain paleoclimatic conditions in accordance with the Eocene global community maturation cycle (Larger Foraminiferal Turnover; Hottinger 1998). The long global warming trend toward the Early Eocene Climate Optimum (EECO; Zachos et al. 2001) with favorable climatic conditions for coralgal reef assemblages persisted, but generally, corals as the main carbonate-producing organisms were absent from this region. In the Adriatic, the shallow-marine carbonate regime is composed of various facies types, which are defined by larger benthic foraminiferal associations and sedimentary structures. These facies are grouped into three main biosedimentary units; BiosZ 2 to BiosZ 4 (Drobne et al. 2007). These zones followed one another in a step-wise geographic pattern and record the temporal and spatial demise of certain ecological conditions. Sedimentation within each zone started with restricted, marginal marine, paralic and pallustrine carbonates that we consider as the initial onset of full marine conditions. Once the marine regime was obtained, the shallow-marine settings supported the development of diverse and abundant foraminiferal assemblages. The biota suggests that the platform jumped into the third Tethyan stage even though sedimentation started in SBZ 1 (Serra-Kiel et

Extended Abstracts

al. 1998) and lasted until SBZ 9 (BiosZ 2), in SBZ 2 until SBZ 12 (BiosZ 3), and in SBZ 11 to SBZ 14 (BiosZ 4). Interestingly, the second Tethyan stage, coralgal facies, is recorded within the NW margin in the BiosZ 2 area only (Zamagini et al. 2009) where the end of the BiosZ 2 coincides with the end of the EECO. The characteristics that differentiate the vast part of the PgAdCP from other Tethyan areas are: •

Reduced foraminiferal species richness in the Paleocene succession (3 assilinid species in Paleocene).

•

Less diverse Eocene foraminiferal assemblages (e.g. 8 species of Nummulites in the Aarly Eocene and 12 species from the Middle Eocene (Pavlovec 2003) vs. 76 for early and 62 for Middle Eocene from the Periediterranean region (Less 1998)).

•

Absence of some taxa (Spiroclypeus sp., Heterostegina sp., Pellatispira sp.).

•

Absence of encrusting foraminifera.

•

The consistent order of occurrence of certain larger foraminiferal groups: complex miliolids and agglutinated conical forminiferas characterize of the beginning of the platform regime, which gave way to alveolinid, orbitolitid and nummulitid assemblages, which eventually were replaced by operculinids and orthophragminids.

•

Concurrence of BiosZ 3 with the Alveolina histrica lineage and Alveolina levantina lineage with BiosZ 4.

We paid special attention to the changes in larger benthic foraminiferal communities during the EECO (between 53–50 Ma, Fig. 1). Within the middle paleolatitude shelf ecosystem and temperatures around 28–32°C (Pearson et al. 2007), larger benthic foraminifera-dominated biota show evidence for typical competitor relations, increased

37


Figure 1 Biostratigraphy and comparison of trends in distribution of larger benthic foraminifera within the Paleogene Adriatic carbonate platform. The shallow benthic zonation was described by Serra-Kiel et al. (1998), and planktonic biostratigraphy (Berggren et al. 1995; Pearson and Berggren 2005), sea-level change curve after Haq and Al-Qahtani (2005). Data from Drobne 1977; Pavlovec 2003; Drobe and Ćosović 2009.

species diversity of alveolinids, and reduced species diversity of larger miliolids. Alveolinids show a great proportion of ecophenotypic variations (dominance of sphaerical, elliptical and fusiform test morphologies). After the EECO, alveolinid diversity drops and the frequency of species with a multichamber proloculus increases (comprising about 15% of total fauna). The post–EECO fauna is characterised by dominance of nummulitids (Nummulites s.str.) that we considered a low metabolic epifauna. The increased numbers of bioeroded ortophragminid and nummulitid tests could be related to the predominance of slow-moving predators or changes in trophic regime. The available habitats for alveolinids were reduced, which could be the result of decreasing temperatures (limited capacity of endosymbionts to thrive in warmer sea-water temperatures), or greater terrestrial runoff that reduced water transparencies and favoured flatter forms (discocylinids and operculinids) to colonize cooler water at the sea-bottom. REFERENCES Berggren, W.A.; Kent, D.V.; Swisher, C.C.; Aubry, M-P. 1995: A revised Cenozoic geochronology and chronostratigraphy. In Berggren, W.A.;

38

Kent, D.V.; Aubry, M-P.; Hardenbol, J. ed. Geochronology, Time Scales and Global Stratigraphic Correlation. SEPM (Society for Sedimentary Geology) Special Publication 54: 129-212. Drobne, K. 1977: Alvéolines paléogènes de la Slovénie et de l'Istrie. Schweizerische Paläontologische Abhandlungen 99: 1-132. Drobne, K.; Ogorelec, B.; Riccamboni, R. 2007: Bangiana hanseni n.gen n.sp. (Forminifera), an index species of Danian age (Lower Paleocene) from the Adriatic carbonate platform (SW Slovenia, NE Italy, Herzegovina). Razprave SAZU 48/1: 5-71. Drobne, K.; Ćosović, V. 2009: Palaeobiogeography of the late Cretaceous to Paleogene Larger Miliolids from tropical to subtropical sea belt (Neotethys to Caribbean). Bulletin Societe Geologique France, in press. Haq, B.U.; Al-Qahtani, A.M. 2005: Phanerozic cycles of sea-level change on the Arabian Platform. GeoArabia 10/2: 127-160. Hottinger, L. 1998: Shallow benthic foraminifera at the Paleocene-Eocene boundary. Strata, Serie 1, 9: 61-64. Less G. 1998: Statistical data of the inner cross protoconch diameter of Nummulites and Assilina from the Schaub Collection. Dela – Opera SAZU 4 razr. 34/2: 183-202.

Extended Abstracts


Pavlovec, R. 2003: The types of nummulitins localities in the Dinarides. RMZ – Materials and Geoenvironments 50: 777-788. Pearson, P.N.; van Dongen, B.E.; Nicholas, C.J.; Pancost, R.D.; Schouten, S.; Singano, J.M.: Wade, B. 2007: Stable warm tropical climate through the Eocene Epoch. Geology 35: 211214. Scheibner, C.; Speijer, R.P. 2008: Late Paleoceneearly Eocene Tethyan carbonate platform evolution – A response to long- and short-term paleoclimatic change. /Earth-Science Reviews/ 90: 71-102 Serra-Kiel, J.; Hottinger, L.; Caus, E.; Drobne, K.; Ferrandez, C.; Jauhri, A.K.; Less, G.; Pavlovec, R.; Pignatti, J.; Samso, J.M.; Schaub, H.; Sirel,

Extended Abstracts

E.; Strougo, A.; Tambaraeu, Y.; Tosquella, J.; Zakrevskaya, E. 1998: Larger foraminiferal biostratigrapy of the Tethyan Paleocene and Eocene. Bulletin Societe Geologique France 169/2: 281-299. Zachos, J.; Pagani, M.; Sloan, L.; Thomas, E.; Billups, K, 2001: Trends, rhythms, and aberrations in global climate 65 Ma to present. Science 292: 686–693. Zamagini, J.: Košir, A.; Mutti, M. 2009: The first microbialite - coral mounds in the Cenozoic (Uppermost Paleocene) from the Northern Tethys (Slovenia): Environmentally-triggered phase shifts preceding the PETM? Palaeogeography Palaeoclimatology Palaeoecology 274: 1-17.

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NEW ZEALAND TERRESTRIAL AND MARGINAL MARINE RECORDS ACROSS THE PALEOCENE–EOCENE TRANSITION E.M. Crouch1, J.I. Raine1, E.M. Kennedy1, L. Handley2 and R.D. Pancost2 1

GNS Science, PO Box 30 368, Lower Hutt: [email protected]; 2Organic Geochemistry Unit, School of Chemistry, University of Bristol, UK.

INTRODUCTION In New Zealand, studies of the Paleocene– Eocene Thermal Maximum (PETM, ca 55 Ma), a time of peak global temperatures and greenhouse conditions, and the Paleocene– Eocene transition have primarily focused on marine successions such as Mead Stream, Marlborough, and Tawanui, Hawkes Bay (e.g. Crouch and Brinkhuis 2005; Hollis et al. 2005). While these sections provide detailed records of changes in the marine realm, terrestrial climate and vegetation changes are less well documented. A significant vegetation change from a gymnosperm-rich Paleocene into Casuarina-dominated, angiosperm-rich, floras of the Early Eocene has previously been documented (Raine 1984; Pocknall 1990), but this does not appear to coincide with PETM warming and the carbon isotope (δ13C) excursion (Crouch and Visscher 2003). Plant micro- and macro-fossil records from Paleocene–Eocene terrestrial and marginal marine sediments at South Island outcrop localities (Mt Somers, Otaio Gorge, midWaipara River) are currently being investigated in conjunction with a drillhole (Kumara-2) from the West Coast (Fig. 1). These studies aim to determine if there are specific vegetation indicators in New Zealand of the PETM, or other Early Eocene hyperthermal events, and to better understand the timing and cause of the notable change in Early Eocene vegetation. KUMARA-2 CORE, WESTLAND The Kumara-2 corehole, located in central Westland, was drilled to a total depth of 1756 m, with the lower 50 m or so passing through Paleocene and Eocene sediments of Brunner and Paparoa coal measures (Carter et al. 1986; Raine 1986). Terrestrial sediments are predominant in the lower 50 m and consist of a mixture of sandstone, mudstones and coals

40

Figure 1 Location of the Kumara-2 core, Mt Somers Quarry and Otaio Gorge section.

(Carter et al. 1986), with total organic contents ranging from 0.5 to 36% (Pancost et al. 2006). Compound specific (n-alkane) δ13C analyses indicate the PETM spans 6–7 m, with a negative shift in δ13C values (ca 4‰) followed by a sharp return to near pre-excursion values (Fig. 2). At the current sampling resolution, the onset of the δ13C excursion appears to be coeval with the beginning of marine sediments, indicating a sea-level rise associated with PETM warming (Sluijs et al. 2008). The marine horizon has abundant Apectodinium dinocysts and the peak abundance of dinocysts occurs just below the most negative δ13C values within the PETM. Pollen of angiosperm taxa associated with thermophilic conditions, such as Cupanieidites orthoteichus and Spinizonocolpites prominatus, are first recorded at the base of the PETM. Several abundance changes in the plant microfossil assemblage, such as an increase in Malvacipollis and Gleicheniidites species and Extended Abstracts


decrease in Proteacidites species, are noted within the PETM (Fig. 2). The plant microfossil and organic biomarker record (e.g. abundance of oleananes (angiosperm biomarkers) increases) both indicate that the PETM was associated with some vegetation change, but currently the timing and precise nature of vegetation change is unclear. Interestingly, the Kumara-2 core records the notable increase in Casuarina pollen (Myricipites harrisii), which has been consistently recognised throughout New Zealand in the Early Eocene and marks a notable long-term shift in vegetation records (e.g. Raine 1984; Pocknall 1990; Crouch and Visscher 2003). The timing of this event is not currently well constrained but may be in the upper part of the New Zealand Waipawan Stage (ca 55.5–53 Ma), perhaps between the short-term PETM and longer-term Early Eocene Climatic Optimum (EECO) warming (Crouch and Brinkhuis 2005). Recent studies have identified additional hyperthermal intervals in the Early Eocene (Lourens et al. 2005; Zachos et al. 2008) and it is possible that the increase in Casuarina pollen is associated with a short-term global warming event. At the current sampling resolution, the increase in

Casuarina pollen is not clearly associated with a negative shift in the δ13C record, a characteristic feature of hyperthermals, and higher resolution analysis is needed. MT SOMERS QUARRY, CANTERBURY Sandy facies of the Broken River Formation dominate sediments at Mt Somers silica sand quarry, Canterbury. Thin silt-clay laminae occur at irregular intervals and where possible these were sampled for palynology. Most samples from the lower half of the ca 110 m section were barren of palynomorphs, although one basal sample yielded a Paleocene palynoflora. A ca 3 m mudstone occurs near the middle of the section and most samples in and above this horizon yielded palynomorphs. Abundant Apectodinium dinocysts are present in the ca 3 m mudstone, indicating an Early Eocene age (Fig. 3). Pollen of angiosperm taxa associated with thermophilic conditions, such as Cupanieidites orthoteichus and Spinizonocolpites prominatus, are also recorded in the mudstone horizon. As in Kumara-2, the first indications of Eocene flora are coincident with the Apectodinium acme and

Figure 2 N-alkane (3 representative records, n-C27–29, shown) organic carbon isotope record from the Kumara-2 core, Westland, along with selected groups of plant microfossils and the marine dinoflagellate cyst genus Apectodinium. The shaded region indicates the negative excursion in the carbon isotope record, interpreted to represent the PETM. Extended Abstracts

41


Figure 3 Stratigraphic record of the Mt Somers Quarry succession. Selected groups of plant microfossils, along with the marine dinoflagellate cyst genus Apectodinium, are shown from the intervals where palynomorphs are preserved.

a marine influence. The mudstone unit may represent the PETM but δ13C analyses have not yet been carried out. The overlying glauconitic sandstone (Homebush Sandstone) appears to span part of the EECO, with marine dinocysts (e.g. Wilsonidium ornatum, Apectodinium spp.) indicating an age within the New Zealand Mangaorapan Stage (ca 53–49.5 Ma). Casuarina pollen (Myricipites harrisii) are a common component of the plant microfossil assemblage (Fig. 3), which again suggests this notable shift in the vegetation record was established by the EECO. OTAIO GORGE, CANTERBURY Coal measures at Otaio Gorge, south Canterbury, contain New Zealand’s earliest known Eocene leaf flora. While the sequence contains coal measures, the presence of Apectodinium and other dinocysts indicates a marine influence through parts of the measured section. Samples were collected for plant macro- and micro-fossil analyses. At present,

42

Paleocene palynomorph assemblages have not been recognised, as the basal part of the section contains poorly preserved palynofloras. It is not clear if the Paleocene–Eocene transition is encompassed in the section, but examination of a leaf flora matrix revealed abundant Apectodinium dinocysts and suggests a possible PETM or earliest Eocene age. Initial analyses of the plant macrofossils indicate a low diversity assemblage, with dominant broadleaf angiosperms. Most dicot angiosperms have smooth margins. A preliminary Leaf Margin Analysis of the plant macrofossil assemblage suggests mean annual temperatures in the mid 20°Cs, although further examination of leaf forms is needed to corroborate this temperature estimate. REFERENCES Carter, M., Kelly, C., Hillyer, M., McDowell, P. 1986: Kumara-2, -2A, well completion report. Petroleum Report Series PR 1183. Ministry of Economic Development. 714 p. Crouch, E.M.; Visscher, H. 2003: Terrestrial vegetation record from across the initial

Extended Abstracts


Eocene thermal maximum at the Tawanui marine section, New Zealand. In Wing, S.; Gingerich, P.; Schmitz, B.; Thomas, E. ed. Causes and Consequences of Globally Warm Climates in the Early Paleogene. Geological Society of America Special Paper 369. Boulder, Colorado, Geological Society of America. Pp. 351-363. Crouch, E.M.; Brinkhuis, H. 2005: Environmental change across the Paleocene-Eocene transition from the eastern margin of New Zealand: a marine palynological approach. Marine Micropaleontology 56: 138-160. Hollis, C.J.; Dickens, G.R.; Field, B.D.; Jones, C.M., Strong, C.P. 2005: The PaleoceneEocene transition at Mead Stream, New Zealand: a southern Pacific record of early Cenozoic global change. Palaeogeography, Palaeoclimatology, Palaeoecology 215: 313343. Lourens, L.; Sluijs, A.; Kroon, D.; Zachos, J.; Thomas, E.; Röhl, U.; Bowles, J.; Raffi, I. 2005: Astronomical pacing of late Paleocene to early Eocene global warming events. Nature 435: 1083-1087. Pancost, R.D.; Handley, L.; Crouch, E.; Hankinson, E.; Steart, D.; Collinson, M.; Pearson, P.; Scott, A. 2006: Using higher plant biomarkers to obtain new carbon isotope records across the PETM. European Geosciences Union (EGU)

Extended Abstracts

General Assembly 2006, Vienna, Austria. Geophysical Research Abstracts 8, 07084, 2006. Pocknall, D.T. 1990: Palynological evidence for the early to middle Eocene vegetation and climate history of New Zealand. Review of Palaeobotany and Palynology 65: 57-69. Raine, J.I. 1984: Outline of a palynological zonation of Cretaceous to Paleogene terrestrial sediments in West Coast region, South Island, New Zealand. New Zealand Geological Survey Report 109. 82p. Raine, J.I. 1986: Palynostratigraphy of Sub-Kaiata Formation Sequence, Petrocorp Kumara-2 well, Westland. New Zealand Geological Survey Report PAL 107. 13p. Sluijs, A.; Brinkhuis, H.; Crouch, E.M.; John, C.M.; Handley, L.; Munsterman, D.; Boharty, S.M.; Zachos, J.C.; Reichart, G.J.; Schouten, S.; Pancost, R.D.; Sinninghe Danste, J.S.; Welters, N.D.; Lotter, A.F.; Dickens, G.R. 2008: Eustatic variations during the PaleoceneEocene greenhouse world. Paleoceanography 23, PA4216, doi:10.1029/2008PA001615. Zachos, J.C.; Dickens, G.R.; Zeebe, R.E. 2008: An early Cenozoic perspective on greenhouse warming and carbon-cycle dynamics. Nature 451: 279-283.

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DYNAMICS OF PLANT – INSECT HERBIVORE INTERACTIONS DURING LATE PALEOCENE AND EARLY EOCENE ENVIRONMENTAL PERTURBATIONS IN THE BIGHORN BASIN, WYOMING, USA E.D. Currano1, C.C. Labandeira2,3 and P. Wilf4 1

Miami University, Department of Geology, OH 45056, USA: [email protected]; 2Smithsonian Institution, Department of Paleobiology, Washington, DC 20013, USA; 3University of Maryland, Department of Entomology, College Park, MD 20742, USA; 4Pennsylvania State University, Department of Geosciences, University Park PA 16082, USA

The Paleocene–Eocene sediments of the Bighorn Basin, Wyoming, USA, record a time of significant variation in climate, biodiversity, and floral and faunal composition. Beginning in the Late Paleocene, global temperatures warmed to the Early Eocene Climatic Optimum (EECO) 51–53 Ma (Zachos et al. 2001). Superimposed on this gradual warming are the abrupt temperature and CO2 increase at the Paleocene–Eocene Thermal Maximum (PETM, 55.8 Ma) and an Early Eocene cool period (Röhl et al. 2000; Wing et al. 2000; Zachos et al. 2003). Although many studies have analyzed the responses of individual taxonomic groups to climate change during this interval (e.g. Chew 2009; Gingerich 2006; Wing 1998), very few have focused on interactions among organisms (Currano et al. 2008). Here, we examine the effect of climate fluctuations that are well documented and

well-placed stratigraphically in the Bighorn Basin on the plant-insect herbivore system there. In particular, we ask whether plants and herbivorous insects responded synchronously to environmental changes. It is important to understand the dynamics of plant–insect interactions because these two groups dominate modern non-microbial terrestrial biodiversity. Today, 70% of herbivorous insect species are oligophagous or monophagous (Bernays and Chapman 1994), making the majority of plant-insect herbivore interactions highly species specific. Therefore, we predict that plants and herbivorous insects responded to climate change synchronously and in the same manner. Alternatively, the relationship between plants and insects could be destabilized by the effects of temperature changes, particularly the abrupt PETM event. A decoupling of this sort has been observed following the CretaceousPaleogene extinction (Wilf et al. 2006).

Figure 1 First and last appearances of plants (A, B) and insect damage morphotypes (C, D) at each stratigraphic level. The data are fully described in the text. First or last appearances cannot be calculated for some intervals, and these are labelled “na.” Intervals with no first or last appearances are labelled “0.”

44

To test these predictions, we examined insect feeding damage on angiosperm leaves from the Bighorn Basin of Wyoming, USA. We conducted insect damage censuses at nine stratigraphic levels ranging in age from 59 to 52.5 Ma (Table 1). These floras occur during the Late Paleocene warming, PETM, Early Eocene cool period, and beginning of the EECO. A total of 9071 fossil angiosperm leaves belonging to 107 species were examined for the presence or absence of 71 insect feeding damage morphotypes (DTs;(Labandeira et al. 2007). Here, we focus on the turnover of plants and DTs as represented by first and last appearances. Turnover Extended Abstracts


of DTs is not directly equivalent to herbivorous insect turnover because some DTs can be made by many different insect groups and some insect groups make multiple DTs. However, our preliminary extant observations (unpublished) show that the DTs on individual host plants correlate at least approximately to real insect species.

PETM flora, which is composed primarily of plants not found at any other time in the Bighorn Basin. A third scenario includes all morphotypes from the nine sites. Except when using the Wing (1998) compilation, the number of first appearances in the Tiffanian 4a flora and last appearances in the Wasatchian 7 flora cannot be determined.

Figure 1 shows the number of first and last appearances of plant morphotypes (A, B) and DTs (C, D) in each flora. Plant stratigraphic ranges were determined using three different methods. First, the range for each morphotype was ascertained using both our data and Wing’s (1998) compilation of plant data for hundreds of localities in the Bighorn Basin; first and last appearances were tabulated accordingly. However, insect damage has been measured only at the nine sites in this study. Therefore, the number of floral first and last appearances was also computed using only the data from the nine sites. In one case, only those morphotypes occurring in more than one flora (boundary-crossers) were included. This scenario minimizes the uniqueness of the

Although many of the major changes in insect damage composition occur at times of major floral turnover, there are subtle differences between the turnover patterns for the two groups. The data show a steady addition of plant morphotypes during the Late Paleocene and a peak in last appearances at the end of the Paleocene (Cf3). Similarly, new DTs appear throughout the Paleocene, but the number of first appearances may be elevated due to edge effects (Foote 2000). Both plant and DT turnover are high during the PETM. This likely represents a northward migration by thermophilic plants and highly specialized insect herbivores in response to warming (Currano et al. 2008; Wing et al. 2005). A peak in floral last occurrences occurs during

USNM Locality Number

Epoch, Mammal Zone

Age (Ma)

Formation

Mean Annual Temperature (oC)

No. of Leaves in census

USNM 42400 - 42406

Eocene, Wasatchian 7

52.75

Willwood

22.2 ± 2 *

1821

37560

Eocene, Wasatchian 5

53.4

Willwood

15.8 ± 2.2 *

693

USNM 37654, 42407, 42408, 42409, 42410

Eocene, Wasatchian 3-4

54.2

Willwood

11.1 ± 2.8║

491

USNM 42395 – 42399

Eocene, Wasatchian 1-2

55.2

Willwood

16.4 ± 2.7 *

1008

USNM 42384

PETM (Eocene), Wasatchian 0

55.8

Willwood

20.1 ± 2.8 ‡

995

USNM 41643

Paleocene, Clarkforkian 3

55.9

Fort Union

16.4 ± 2.9║

843

USNM 42411

Paleocene, Clarkforkian 2

56.4

Fort Union

12 ± 3║

1016

USNM 42042

Paleocene, Tiffanian 5b

57.5

Fort Union

10.5 ± 2.9 †

1364

USNM 42041

Paleocene, Tiffanian 4a

58.9

Fort Union

10.5 ± 2.9 †

840

Table 1 Sampling summary. The Wasatchian 7 and Wasatchian 1-2 levels have more than one USNM locality number because multiple quarries were dug within the same bed. Leaf fossils from the cool period (Wasatchian 3-4) are scarce, and those included here come from four quarries in the same bed (USNM 4240742410) and an additional quarry that is 8.5 miles away but within 5 meters stratigraphically (USNM 37654). Temperature estimates are from the following sources: * Wing et al. (2000) Extended Abstracts

║Currano (2008)

‡ Wing et al. (2006)

† Currano et al. (2008)

45


the cool period (Wa2/3), followed by an increase in floral first appearances during the EECO (Wa7; Wing 1998). The insect damage data show a similar pattern, although there are more last appearances in Wa5 than Wa2/3, perhaps suggesting a lag between changes in floral and insect damage composition. This implies that plants and insect herbivores do not always respond synchronously to climate change. ACKNOWLEDGEMENTS We thank Scott Wing for his assistance on all aspects of this project and Liz Lovelock for morphotyping the PETM flora. Additionally, we thank the many students and colleagues who assisted us in the field. This work was supported by an NSF Graduate Research Fellowship, Petroleum Research Fund grant 40546-AC8, the Roland Brown Fund, and student research grants from the Evolving Earth Foundation, GSA, the Paleontological Society, and Pennsylvania State University. REFERENCES Bernays, E.A.; Chapman, R.F. 1994: Host-Plant Selection by Phytophagous Insects. London, Chapman and Hall. 312 p. Chew, A.E. 2009: Paleoecology of the early Eocene Willwood mammal fauna from the central Bighorn Basin, Wyoming. Paleobiology 35: 13-31. Currano, E.D. 2008: Variations in insect herbivory on angiosperm leaves through the late Paleocene and early Eocene in the Bighorn Basin, Wyoming, USA. PhD thesis. Pennsylvania State University, State College, PA. Currano, E.D.; Wilf, P.; Wing, S.L.; Labandeira, C.C.; Lovelock, E.C.; Royer, D.L. 2008: Sharply increased insect herbivory during the Paleocene-Eocene Thermal Maximum. Proceedings of the National Academy of Sciences 105(6): 1960-1964. Foote, M. 2000: Origination and extinction components of taxonomic diversity: general problems. Paleobiology 26(Suppl. to No. 4): 74-102.

46

Gingerich, P.D. 2006: Environment and evolution through the Paleocene-Eocene thermal maximum. Trends in Ecology and Evolution 21: 246-253. Labandeira, C.C.; Wilf, P.; Johnson, K.R.; Marsh, F. 2007: Guide to Insect (and Other) Damage Types on Compressed Plant Fossils. Version 3.0. Washington, D.C., Smithsonian Institution. 25 p. Röhl, U.; Bralower, T.J.; Norris, R.D.; Wefer, G. 2000: New chronology for the late Paleocene thermal maximum and its environmental implications. Geology 28: 927-930. Wilf, P.; Labandeira, C.C.; Johnson, K.R.; Ellis, B. 2006: Decoupled plant and insect diversity after the end-Cretaceous extinction. Science 313: 1112-1115. Wing, S.L. 1998: Late Paleocene-Early Eocene floral and climatic change in the Bighorn Basin, Wyoming. In Aubry, M. P.; Lucas, S. G.; Berggren, W. A., ed. Late Paleocene Early Eocene climatic and biotic events in the marine and terrestrial records. New York, Columbia University Press. Pp. 380-400. Wing, S.L.; Bao, H.; Koch, P.L. 2000: An early Eocene cool period? Evidence for continental cooling during the warmest part of the Cenozoic. In Huber, B. T.; MacLeod, K. G.; Wing, S. L., ed. Warm climates in earth history. Cambridge, Oxford University Press. Pp. 197-237. Wing, S.L.; Harrington, G.J.; Smith, F.A.; Bloch, J.I.; Boyer, D.M.; Freeman, K.H. 2005: Transient floral change and rapid global warming at the Paleocene-Eocene boundary. Science 310: 993-996. Wing, S.L.; Lovelock, E.C.; Currano, E.D. 2006. Climatic and floral change during the PETM in the Bighorn Basin, Wyoming, USA. Climate and Biota of the Early Paleogene. 2006. Volume Abstracts. Bilbao, 168p Zachos, J.; Pagani, M.; Sloan, L.; Thomas, E.; Billups, K. 2001: Trends, rhythms, and aberrations in global climate 65 Ma to present. Science 292: 686-693. Zachos, J.C.; Wara, M.W.; Bohaty, S.; Delaney, M.L.; Petrizzo, M.R.; Brill, A.; Bralower, T.J.; Premoli-Silva, I. 2003: A transient rise in tropical sea surface temperature during the Paleocene-Eocene Thermal Maximum. Science 302: 1551-1554.

Extended Abstracts


ELEMENTAL GEOCHEMISTRY AND DIAGENESIS IN EOCENE TROPICAL PLANKTONIC FORAMINIFERA: INSIGHTS FROM LASER–ABLATION PROFILING Andrea Dutton1, Paul Pearson2, Tim Bralower3, Bridget Wade4, Stephen Eggins1 and Claire Thompson1 1

2

Research School of Earth Sciences, Australian National University, Australia: [email protected]; School of Earth, Ocean and Planetary Sciences, Cardiff University, Cardiff, UK; 3Department of Geosciences, Pennsylvania State University, University Park PA, USA; 4Department of Geology and Geophysics, Texas A&M, College Station, TX, USA.

In the past few decades, a wealth of new multiproxy climate data has become available for describing conditions in the Eocene greenhouse world. A reoccurring theme in these studies is to resolve a discrepancy that has long been recognized between tropical temperatures predicted by models and those reconstructed from planktonic foraminiferal oxygen isotope compositions (δ18O) (Huber 2008). The cooler (near modern) tropical sea surface temperatures (SSTs) reconstructed from planktonic δ18O recovered in deep sea cores are at odds with considerably warmer tropical SSTs predicted by models. Most arguments regarding this cool tropics paradox invoke extensive diagenesis of planktonic foraminifera or fundamental gaps in our understanding of climate dynamics as possible explanations. Unfortunately, decisive evidence for diagenetic alteration of the chemical signals carried by these shells has been hard to come by. In the past several years, exceptionally wellpreserved “glassy” foraminifera of Eocene age have been recovered from clay-rich hemipelagic sections (Pearson et al. 2001, 2007). These glassy foraminifera also have oxygen isotope compositions that are suggestive of warmer tropical SSTs — more in keeping with those predicted by the models. Specifically, the δ18O values of planktonic foraminifera from pelagic sites are significantly more positive (by up to 2‰) suggesting cooler temperatures (by ~10ºC) than the hemipelagic glassy foraminifera. The striking difference in calcite textures between these glassy foraminifera and those recovered from carbonate-rich pelagic sections is clear evidence of different post-depositional preservation in these environments, though it is less clear how this process affects foraminiferal chemistry.

Extended Abstracts

We are using laser ablation inductivelycoupled mass spectrometry (LA-ICP-MS) to examine changes in elemental geochemistry across the thickness of the foraminiferal shell walls in tropical Eocene planktonic foraminifera with different preservation states to better understand the geochemical signal of the altered (diagenetic) component. These laser ablation profiles enable us to detect micronscale diagenetic calcite, to calculate the thickness of alteration rims, and ultimately to evaluate how extensively the primary signal of the entire shell has been altered. If a dual composition layering is present (i.e. secondary inorganic calcite encrusting primary foraminiferal calcite) the unique high spatial resolution of this laser system is capable of resolving the elemental compositions of these separate components. In most of the Eocene-age specimens studied from Shatsky Rise (ODP Site 1209) and Alison Guyot (ODP Site 865), the outer edge of the shell is characterized by higher Mg/Ca ratios, lower Sr/Ca ratios, and elevated concentrations of other trace metal cations that suggest the presence of some secondary calcite (Fig. 1). While the Sr/Ca ratios of this secondary calcite are within the range of predicted values according to ‘typical’ laboratory-derived distribution coefficients for inorganic calcite (Banner 1995; Holland 1966) and measured pore water values (Bralower et al. 2003; Sager et al. 1993), the Mg/Ca ratios are only slightly elevated along the outer shell margin, typically by about 1 mmol/mol. This results in absolute compositions that are almost an order of magnitude lower than those predicted using laboratory-derived distribution coefficients (e.g. Mucci and Morse 1983). This finding is consistent with the results obtained by Sexton et al. (2006) who compared Mg/Ca ratios of foraminifera from carbonate-rich deep-sea

47


Figure 1 Scanning electron microscope (SEM) image of bladed secondary calcite that nucleates on topographically prominent muricae. Scale bar is 10 microns.

cores to those of glassy foraminifera recovered from clay-rich sections and found only a slight elevation in the Mg/Ca ratio of the deep-sea specimens. We propose two possible explanations for this observation: •

the first possibility is in keeping with the conclusion of Sexton et al. (2006) that the distribution coefficient of Mg into inorganic calcite in deep-sea sediments is more consistent with those suggested by Baker et al. (1982) for this environment rather than those derived from laboratory experiments (e.g. Mucci and Morse 1983); or

•

that the amount of secondary calcite is so minimal in volume that the Mg/Ca signal of this component is swamped by the composition of the primary shell wall.

The second of these hypotheses is potentially consistent with the observation of conspicuous

48

secondary calcite crystals that nucleate on muricate structures (Fig. 2) and comprise a minimal contribution to the overall volume of foraminiferal carbonate — assuming that these crystals are secondary and the remainder of the material is primary in composition. We are currently testing these two hypotheses using maps of Mg/Ca for cross-sections of chamber walls using an electron microprobe, similar in technique to Eggins et al. (2004). This technique offers another way to quantify the Mg/Ca composition across the shell wall at micron-scale intervals with the advantage that the mapped image provides unambiguous association between Mg/Ca values and specific parts of the foraminiferal test. In comparison to the laser ablation technique, it should be emphasized the electron microprobe technique requires two orders of magnitude more time than LA-ICP-MS, and also that most of the additional elements that we are able to measure with LA-ICP-MS (Sr, Al, Fe, Zn, Ba, etc.) are near or below detection limits on the electron microprobe. Hence this method is a useful, complementary tool to the laser ablation profiles, but does not provide the quantity and quality of data that can be resolved using LAICP-MS. Using this combination of laser ablation and electron microprobe data, we will be able to quantitatively discuss the geochemical evidence for the presence and degree of alteration in tropical planktonic foraminifera

Figure 2 Mg/Ca and Sr/Ca profiles for four chambers of an individual Morozovella sp. specimen from Alison Guyot (865C-8H-6, 10-11). Note high Mg/Ca and low Sr/Ca values on the very outer margin of the chamber wall which are commonly associated with enrichment in other trace metals (e.g. Zn, Mn, Fe, Ba). Final chamber (f), penultimate chamber (f-1), etc. Extended Abstracts


recovered from pelagic carbonate-rich sediments in comparison to glassy foraminifera recovered from clay-rich hemipelagic sites. Ultimately, we are also interested in drawing comparisons between tropical and high-latitude sites and between specimens from the warm Early Eocene and from the Mid- to Late Eocene when cooler bottom waters may change the degree and even the mechanism of diagenesis in deep-sea sediments. REFERENCES Baker, P.A.; Geiskes, J.M.; Elderfield, H. 1982: Diagenesis of carbonates in deep-sea sediments--Evidence from Sr/Ca ratios and interstitial dissolved Sr2+ Data. Journal of Sedimentary Petrology 52: 71-82. Banner, J. 1995: Application of the trace element and isotope geochemistry of strontium to studies of carbonate diagenesis. Sedimentology 42: 805-824. Bralower, T.J.; Premoli Silva, I.; Malone, M.J.; Arthur, M.A.; Averyt, K.; Bown, P.R.; Brassell, S.C.; Channell, J.E.T.; Clarke, L.J.; Dutton, A. 2003. Proceedings of the Ocean Drilling Program, Initial Reports, Vol. 198, College Station, TX. Eggins, S.M.; Sadekov, A.; De Deckker, P. 2004: Modulation and daily banding of Mg/Ca in Orbulina universa tests by symbiont photosynthesis and respiration: a complication for seawater thermometry? Earth and Planetary Science Letters 225: 411-419.

Extended Abstracts

Holland, H.D. 1966. The coprecipitation of metallic ions with calcium carbonate. Princeton University Annual Report Princeton University, Department of Geology. Contribution No. AT (30-I)-2266. Huber, M. 2008: A Hotter Greenhouse? Science 321: 353-354. Mucci, A.; Morse, J. 1983: The incorporation of Mg2+ and Sr2+ into calcite overgrowths: Influences of growth rate and solution composition. Geochimica et Cosmochimica Acta 47: 217-233. Pearson, P.N.; Ditchfield, P.; Singano, J.; HarcourtBrown, K.; Nicholas, C.; Olsson, R.; Shackleton, N.J.; Hall, M.A. 2001: Warm tropical sea surface temperatures in the late Cretaceous and Eocene epochs. Nature 413: 481-485. Pearson, P.N.; van Dongen, B.E.; Nicholas, C.J.; Pancost, R.D.; Schouten, S.; Singano, J.; Wade, B.S. 2007: Stable warm tropical climate through the Eocene Epoch. Geology 35: 211214. Sager, W.W.; Winterer, E.L.; Firth, J.V.; et al. 1993. Proceedings of the Ocean Drilling Program, Initial Reports. Ocean Drilling Program, College Station, TX. Sexton, P.F.; Wilson, P.A.; Pearson, P.N. 2006: Microstrucural and geochemical perspectives on planktic foraminiferal preservation: "Glassy" versus "Frosty". Geochemistry, Geophysics, Geosystems 7: Q12P19.

49


PALEOGENE DEEP-WATER DEPOSITS AT GAMS (AUSTRIA): FROM THE K/PG BOUNDARY TO THE P/E BOUNDARY IN A TETHYAN SETTING H. Egger1, C. Koeberl2, C. Spötl3, M. Wagreich2 and O. Mohamed4 1

Geological Survey of Austria, Neulinggasse 38, 1030 Vienna, Austria: [email protected]; 2University of Vienna, Althanstrasse 14, 1090 Vienna, Austria; 3University of Innsbruck, Innrain 52, 6020 Innsbruck, Austria; 4University of Graz, Heinrichstrasse 26, 8010 Graz.

Stratigraphical and sedimentological investigations have been carried out on the Paleogene deposits of the Gosau Group near Gams (Austria), 120 km southwest of Vienna (Fig. 1). The ca 400 m thick sedimentary succession was deposited in a middle to lower bathyal environment and comprises the Paleocene to lowermost Eocene (Fig. 2). The Cretaceous–Paleogene (K/Pg) boundary has been recognized in the newly discovered Gamsbach section (Fig. 3), which encompasses the upper part of the Cretaceous Nephrolithus frequens Zone (CC26) and the lower part of the Paleocene Markalius inversus Zone (NP1). The boundary is characterized by •

an enrichment of the contents of the siderophile elements Ir (6ppb), Co (56ppm), Ni (80ppm), and Cr (130ppm) compared to background and continental crustal values,

•

a sudden decrease of carbon and oxygen isotopic ratios,

•

a sudden decrease of carbonate content from an average of 60wt% to 20wt% , and

•

an acme of the calcareous dinoflagellate cyst Operculodinella operculata, which is succeeded by an acme of the small coccolith species Neobiscutum parvulum.

Figure 1

50

The presence of the Ir-rich boundary layer and the distribution patterns of calcareous nannoplankton indicate a complete sedimentary record across the Cretaceous– Paleogene transition at the Gamsbach section. Carbonate contents and carbon isotope values declined at the K–Pg boundary and did not recover in the Paleogene part of the section, which encompasses the lower half of Zone NP1. The preservation of very small nannoplankton species like Neobiscutum parvulum indicates that the decrease in carbonate was not an effect of increased carbonate dissolution but of lower productivity of calcareous plankton. Additionally, enhanced terrestrially derived input might have diluted the calcareous deposits. Enhanced continental erosion is implied from the reworking of Campanian strata at and above the K–Pg boundary and might have been the result of a substantial sea-level drop at the end of the Maastrichtian. The lithological change at the K–Pg boundary is associated with an abrupt negative shift of bulk stable isotope values. The high degree of co-variation between δ13C and δ18O in the Gamsbach section implies diagenetic overprint. SEM studies show that calcite overgrowths on coccoliths are more extensive in Maastrichtian than in earliest Paleocene

Occurrences of the Gosau Group in Austria. Extended Abstracts


Figure 2 Stratigraphic and lithologic log of Paleogene part of the Gosau Group at Gams, including bulk stable isotope values and the occurrence of Apectodinium augustum.

samples. The latter show abundant clay flakes and scattered, rather well preserved nannofossils. In case part of the significant drop in δ18O at the K/Pg boundary indeed reflects a pristine marine signal (which will be

Extended Abstracts

difficult to verify unambiguously) the data suggest a rapid warming of the photic zone (where most of the plankton lived) followed by a transient cooling of ocean surface waters.

51


Figure 3

The Gamsbach section (the person is sitting on the top of the Maastrichtian).

The decrease in ocean surface water temperatures was probably not an effect of upwelling, because eutrophic species such as Braarudosphaera bigelowii are very rare at the Gamsbach section. Instead, the cooling event is associated with a bloom of Neobiscutum parvulum, whereas coevally the abundance of Operculodinella operculata decreases. This suggests that water temperature was a limiting factor for the mass-occurrences of both species. In different areas these blooms occur at different stratigraphic levels. This may

Figure 4 Apectodinium augustum from sample 11/08 at the Gams section.

52

reflect a number of short-lived changes in the configuration of ocean circulation after the impact. In the basal Danian of the Gamsbach section, five nannoplankton species have their first occurrences (FO). These new species are Cyclagelosphaera alta (FO +2 cm), Markalius astroporus (FO +4 cm), Lanternithus duocavus (FO +12 cm), Biantholithus sparsus (FO +30 cm), and Markalius apertus (FO +100 cm). The Danian deposits are characterized by a predominance of red and grey pelagic to hemipelagic marlstones and marly limestones. Thin sandstone-calcarenite turbidite beds are present in variable amounts, but sandstone to pelite ratios stay below 1:5. Turbidite beds are typically calcarenitic with 37 Ma. Crown-Cetacea, Neoceti, evolved from amongst such basilosaurid archaeocetes.

Extended Abstracts


Surprisingly, archaeocete-like Cetacea persisted after the Neoceti appeared, with significant Late Oligocene (~26 Ma) records from New Zealand (Fordyce 2004). These latesurviving archaeocetes coexisted with filterfeeding mysticetes and echolocating dolphins, which raises questions about ecological partitioning in Oligocene Cetacea. To consider the Neoceti, or crown-Cetacea, the two major living clades are well represented by Oligocene fossils of Mysticeti - baleen whales (filter-feeders) and relatives, and Odontoceti toothed whales and dolphins (echolocators). The current oldest named species in these groups are: Llanocetus denticrenatus (Mysticeti) from Antarctica, about 34.2 Ma (latest Eocene), and Simocetus rayi (Odontoceti) from northeastern Pacific, about 32 Ma. The initial appearance and radiation of odontocetes and mysticetes coincides with the opening of the Southern Ocean and change from greenhouse to icehouse earth, suggesting that tectonic (physical, “bottom up”) processes drove major aspects of cetacean evolution through influence on food supplies (Fordyce 1980; see also Lindberg and Pyenson 2007, Berger 2007). There is no evidence that Eocene–Oligocene cooling had any influence on the evolution of cetacean brain size (Marino et al. 2007). The diversity of Late Oligocene clades implies major radiation and ecological partitioning before 30 Ma (Fordyce 2003), an idea supported by molecular phylogenies (e.g. Nikaido et al. 2001). Molecules, however, do not capture the radiations and diversity of Oligocene groups that are now extinct, such as Squalodontidae, Waipatiidae, and Squalodelphinidae — all families of platanistoid dolphins. This reveals a clear limit to the application of molecular approaches to “deep time” patterns and processes. PENGUINS Penguins represent an exclusively southern clade of wing-propelled diving seabirds which cannot fly in air. They are similar in body form and habits to some Northern Hemisphere auks (Alcidae) and extinct northern pelicanrelatives, the plotopterids. Origins probably lie with tube-nosed birds, Procellariformes (petrels, albatrosses). For many decades, the oldest fossils were Late Eocene species, but

Extended Abstracts

named species (of Waimanu, below) are now known from the middle of the Paleocene. The Paleogene record of penguins is patchy, with significant records from New Zealand, Antarctica, and South America — most recently including unexpected fossils from Peru (e.g. Fordyce and Jones 1990, Clarke et al. 2007, Jadwiszczak 2009). Single fossil bones are much more common than partial skeletons and, regrettably, single bones are the basis for naming many species. For example, the first fossil penguin named, Palaeeudyptes antarcticus, is known only from an incomplete ankle bone of Late Eocene or Early Oligocene age, from Kakanui, New Zealand. The specimen clearly represents a large extinct penguin, and was new to science when Huxley (1859) named it, but the sole known bone is a poor basis on which to compare new material. Surprisingly, incomplete single bones continue to be used as type specimens (e.g. Tambussi et al. 2005), in spite of the increasing record of partial (or even reasonably complete) skeletons. Paleocene fossils of Waimanu from New Zealand significantly extend the record of named penguins to about 58–61.5 Ma, and reveal skeletal form and lifestyle early in penguin history (Fordyce and Jones 1990, Slack et al. 2006). Waimanu helps calibrate the molecular clock of crown-birds (Neornithes), and points to a significant bird radiation in the late Cretaceous (Slack et al. 2006). Feduccia (2003), however, portrayed the radiation of Neornithes as a Cenozoic event, stemming from one or a few bird lineages that survived the K–T boundary. New fossils from New Zealand, South America and Antarctica (Clarke et al. 2007, Ando 2007, Jadwiszczak 2009) continue to fill the Paleocene–Eocene gap. Baker et al. (2006) used Waimanu to calibrate a molecular phylogeny of penguins, allowing the prediction that the lineages of modern Antarctic penguins perhaps arose ca 40 Ma, associated with Antarctic cooling. However, all the known Paleogene penguins are stem-taxa (Clarke et al. 2007), which originated before the most recent common ancestor of extant species, and there is no hint of crown-taxa from the Paleogene (including the well sampled New Zealand Late Oligocene; Ando 2007). Fossils in the crownpenguin genera are all late Neogene, no older

61


than about 12 Ma and, judging from uncertainty about the ages of key New Zealand species, possibly as young as 6 Ma. The stratigraphically-calibrated cladogram of Clarke et al. (2007) shows that penguins clearly survived the end–Paleocene thermal event; otherwise, there appears to be no strong correlation of Paleogene diversity or cladogenesis either with warming or cooling. Eocene and Oligocene penguins include species significantly larger than living forms, with inferred body masses >60 kg; since these lived at times of warmth, there can be no serious correlation of large size with cool climates. A few fragments caution that not all Paleogene penguins were big; there were also diminutive species, as in modern seas. REFERENCES Ando, T. 2007: Functional change in penguin phylogeny. Ph.D. thesis: Dunedin, The Library, University of Otago. Bajpai, S.; Gingerich, P.D. 1998: A new Eocene archaeocete (Mammalia, Cetacea) from India and the time of origin of whales. Proceedings of the National Academy of Sciences of the United States of America 95: 15464-15468. Baker, A.J.; et al. 2006: Multiple gene evidence for expansion of extant penguins out of Antarctica due to global cooling. Proceedings of the Royal Society B - Biological Sciences 273: 11-17. Barnosky, A.D. 2001: Distinguishing the effects of the Red Queen and Court Jester on Miocene mammal evolution in the northern Rocky Mountains. Journal of Vertebrate Paleontology 21: 172-185. Berger, W.H. 2007: Cenozoic cooling, Antarctic nutrient pump, and the evolution of whales. Deep Sea Research II 54: 2399-2421. Clarke, J.A.; et al. 2007: Paleogene equatorial penguins challenge the proposed relationship between biogeography, diversity, and Cenozoic climate change. Proceedings of the National Academy of Sciences of the United States of America 104: 11545-11550. Feduccia, A. 2003: 'Big bang' for tertiary birds? Trends in Ecology and Evolution 18: 172-176. Fordyce, R.E. 1980: Whale evolution and Oligocene Southern Ocean environments. Palaeogeography, Palaeoclimatology, Palaeoecology 31: 319-336. Fordyce, R.E. 2003: Cetacean evolution and Eocene-Oligocene oceans revisited. In Prothero, D.R.; et al. ed. From Greenhouse to

62

Icehouse: The marine Eocene-Oligocene transition. New York, Columbia University Press. Pp. 154-170. Fordyce, R.E. 2004: The transition from Archaeoceti to Neoceti: Oligocene archaeocetes in the Southwest Pacific. Journal of Vertebrate Paleontology 24: 59A. Fordyce, R.E.; Jones, C.M. 1990: The history of penguins, and new fossil penguin material from New Zealand. In Davis, L.S.; Darby, J.D. ed. Penguin biology. San Diego, Academic Press. Pp. 419-446. Geisler, J.H.; Theodor, J.M. 2009: Hippopotamus and whale phylogeny. Nature 458: E1-E4. Gingerich, P.D. 2005: Cetacea. In Rose, K.D.; Archibald, J.D. ed. Placental mammals: origin, timing, and relationships of the major extant clades. Baltimore, Johns Hopkins University Press. Pp. 234-252. Gingerich, P.D. 2006: Environment and evolution through the Paleocene-Eocene thermal maximum. Trends in Ecology and Evolution 21: 246-253. Gingerich, P.D.; et al. 2009: New protocetid whale from the middle Eocene of Pakistan: birth on land, precocial development, and sexual dimorphism. PLoS ONE 4: e4366 doi:10.1371/journal.pone.0004366. Huxley, T.H. 1859: On a fossil bird and a fossil cetacean from New Zealand. Quarterly Journal of the Geological Society of London 15: 670677. Jadwiszczak, P. 2009: Penguin past: The current state of knowledge. Polish Polar Research 30: 3-28. Lindberg, D.R.; Pyenson, N.D. 2007: Things that go bump in the night: evolutionary between cephalopods and cetaceans in the Tertiary. Lethaia 40: 335-343. Marino, L.; et al. 2007: Cetaceans have complex brains for complex cognition. Plos Biology 5: 966-972. Maxwell, P.A. 1992: Eocene Mollusca from the vicinity of McCulloch's Bridge, Waihao River, South Canterbury, New Zealand: paleoecology and systematics. New Zealand Geological Survey Paleontological Bulletin 65: 1-280. Nikaido, M.; et al. 2001: Retroposon analysis of major cetacean lineages: The monophyly of toothed whales and the paraphyly of river dolphins. Proceedings of the National Academy of Sciences of the United States of America 98: 7384-7389.

Extended Abstracts


Rose, K.D. 2006: The beginning of the age of mammals. Baltimore, Johns Hopkins University Press. 448 p. Slack, K.E.; et al. 2006: Early penguin fossils, plus mitochondrial genomes, calibrate avian evolution. Molecular Biology and Evolution 23: 1144-1155. Tambussi, C.P.; et al. 2005: Crossvallia unienwillia, a new Spheniscidae (Sphenisciformes, Aves) from the Late Paleocene of Antarctica. Geobios 38: 667-675. Thewissen, J.G.M.; et al. 2007: Whales originated from aquatic artiodactyls in the Eocene epoch of India. Nature 450: 1190-1194. Uhen, M.D. 2008: New protocetid whales from Alabama and Mississippi, and a new cetacean

Extended Abstracts

clade, Pelagiceti. Journal Paleontology 28: 589-593.

of

Vertebrate

Uhen, M.D.; Pyenson, N.D. 2007: Diversity estimates, biases, and historiographic effects: resolving cetacean diversity the Tertiary. Palaeontological Electronica 10: 11A (1-22). Woodburne, M.O.; Case, J.A. 1996: Dispersal, vicariance, and the Late Cretaceous to Early Tertiary land mammal biogeography from South America to Australia. Journal of Mammalian Evolution 3: 121-161. Worthy, T.H.; et al. 2006: Miocene mammal reveals a Mesozoic ghost lineage on insular New Zealand, southwest Pacific. Proceedings of the National Academy of Sciences of the United States of America 103: 19419-19423.

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MAASTRICHTIAN–DANIAN BENTHIC FORAMINIFERIDS FROM WADI NUKHAL, SINAI, EGYPT: AN INTEGRATED BIOZONATION G. Galal1 and S. Kamel2 1

Geology Department, Faculty of Sciences, Alexandria University, Alexandria, Egypt; 2High Institute Arts, King Mariout, Alexandria, Egypt.

INTRODUCTION Late Cretaceous–Early Paleogene rocks in Egypt are well known for their classical stratigraphic importance. Such rocks provide abundant evidence of several significant events during this interval of geologic time (e.g. the Dababiya GSSP, Ouda and Berggren 2003). The section of Wadi Nukhul, West Central Sinai, Egypt (33º 15’ 23”E, 29º 01’ 47”N) is described here as one of the valuable successions that allowed one of us (G.G.) to establish three biozonal schemes based on bolivinoides, spiral planktic and serial planktic forms (Galal 2003). This abstract introduces additional benthic foraminifera biozonal schemes based on taxa from five suborders, including Astrorhizina, Haplophragmiina, Textulariina, Lagenina and Rotalliina, from the same succession. This was done by grouping related taxa on separate charts and then examining the ranges of short-ranging taxa to establish or recognize biozones based on their ranges. BIOSTRATIGRAPHY The initial use of foraminiferids in biostratigraphy started in the 19th century and emphasized benthic forams, but present day work depends more upon planktics. This may be explained by the role of planktic forams in biozonations where they have been considered more reliable. However, recent attention is again focusing on the use of benthics for biostratigraphic refinement, particularly after the pioneer work of Bolli et al. (1994). A glance at distribution charts of benthic forams in published papers often suggests such taxa have biostratigraphic value. 139 species of benthic foraminiferids, belonging to 76 genera (excluding 7 species related to genus Bolivinoides, which have been previously studied before by G.G.), are recorded from 53 samples collected from the 22.3 m thick Maastrichtian–Danian interval of 64

Wadi Nukhul section, West Central Sinai, Egypt. Ranges of the recorded taxa were plotted, with 74 species (belonging to 58 genera) considered to be long ranging (i.e. recorded from all samples and some of these taxa were found in the literature to begin earlier than the Maastrichtian and to continue after the Danian), while the remaining 65 species (belonging to 44 genera) were found to be short ranging (i.e. range within the studied interval). Haplophragmiine Zonal Scheme Consideration of ranges of the 15 short ranging agglutinated foraminifera species, of 14 genera (out of 34 species of 22 genera), led to the subdivision of the studied interval into 5 Haplophragmiine biozones (based on ranges of 6 species belonging to 6 genera; Fig.1) as follows, from base to top: 1) The Repmanina favilla-Glomospira confusa Interval Zone (Early to Late Maastrichtian). 2) The Glomospira confusa-Tritaxia barakai Concurrent Zone (Late Maastrichtian). 3) The Tritaxia barakai-Kolchidina paleocenicus Interval Zone (Early Danian). 4) The Kolchidina paleocenicus-Vulvulina colei Interval Zone (Middle Danian). 5) The Vulvulina colei Interval Zone (Late Danian). Nodosariacean Zonal Scheme Consideration of ranges of the15 short ranging Lagenine species, of 10 genera (out of 46 species of 22 genera), led to the subdivision of the studied interval into 5 Nodosariacean biozones based on ranges of 4 Nodosariacean species (Fig.1) as follows, from base to top: 1) The Saracenaria navicula Interval Zone (Early to Late Maastrichtian).

Extended Abstracts


Figure 1 Ranges of marker species of the proposed benthic foramininferal biozones in correlation with spiral planktic biozones of Galal (2003) recorded from the Maastrichtian-Danian rocks of Wadi Nukhul, West Central Sinai, Egpyt.

2) The Saracenaria navicula-Neoflabellina delicatissma Interval Zone (Late Maastrichtian)

Coryphostoma biozones based on ranges of 3 Coryphostoma species and 2 Bulimina species (Fig.1) as follows, from base to top:

3) The Neoflabellina delicatissmaNeoflabellina esnehensis Interval Zone (Early Danian).

1) The Coryphostoma incrassata Interval Zone (Early to Late Maastrichtian).

4) The Neoflabellina esnehensis-Dentalina rancocasensis Interval Zone (Middle Danian). 5) The Dentalina rancocasensis Zone (Late Danian).

Interval

Serial Rotaliine Zonal Scheme Consideration of ranges of the 16 short ranging Serial Rotaliine species, of 9 genera (out of 25 species of 15 genera), led to the subdivision of the studied interval into 6 Bulimina and Extended Abstracts

2) The Coryphostoma Coryphostoma crassum (Late Maastrichtian)

incrassataInterval Zone

3) The Coryphostoma crassum-Bulimina quadrata Interval Zone (Early Danian). 4) The Bulimina quadrata-Bulimina asperoaculiata Interval Zone (Early Danian). 5) The Bulimina asperoaculiataCoryphostoma midwayensis Interval Zone (Middle Danian)

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6) The Coryphostoma midwayensis Interval Zone (Late Danian).

6) The Gavelinella lellingensis-Valvulineria aegyptiaca Interval Zone (Middle Danian).

Spiral Rotaliine “Chilostomellacean” Zonal Scheme

7) The Valvulineria aegyptiacaAnomalinoides aegyptiacus Interval Zone (Late Danian).

Consideration of ranges of the 19 short ranging Spiral Rotaliine species, of 11 genera (out of 34 species of 17 genera), led to the subdivision of the studied interval into 8 Chilostomellacean biozones based on ranges of 8 Chilostomellacean species (Fig.1) as follows, from base to top: 1) The Stensioeina excolata-Anomalinoides simplix Interval Zone (E.-Late Maastrichtian). 2) The Anomalinoides simplix-Stensioeina excolata Concurrent Zone (Late Maastrichtian). 3) The Stensioeina excolata-Anomalinoides granosus Interval Zone (Early Danian). 4) The Anomalinoides granosusParalabamina lunata Interval Zone (Early Danian).

8) The Anomalinoides aegyptiacus Interval Zone (Late Danian). REFERENCES Bolli, H.M.; Beckmann, J.P.; Saunders, J.B. 1994: Benthic foraminiferal biostratigraphy of the south Caribbean region. Cambridge University Press. 408p. Galal, G. 2003: Bolivinoid and planktic foraminiferal biostratigraphy along the Cretaceous/Paleogene boundary from Wadi Nukhul, west Cetral Sinai, Egypt. 3rd International Conference on the Geology of Africa, Assiut University, Egypt: 745-776. Ouda, K.; Berggren, W.A. 2003: Biostratigraphic correlation of the Upper Palecene – Lower Eocene succession in the Upper Nile Valley: Asynthesis. Micropaleontology 49, suppl.1: 179-212.A variety of PETM RECORDS in different settings, northeastern Peri-Tethys

5) The Paralabamina lunata-Gavelinella lellingensis Interval Zone (Middle Danian).

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Extended Abstracts


A VARIETY OF PETM RECORDS IN DIFFERENT SETTINGS, NORTHEASTERN PERI-TETHYS Y. Gavrilov, E. Shcherbinina, O. Golovanova and B. Pokrovsky Geological Institute of Russian Academy of Sciences, Pyzhevsky 7, Moscow 119017, Russia: [email protected]

INTRODUCTION

RESULTS

The impact of the Paleocene–Eocene Thermal Lithological and organic matter Maximum (PETM) is clearly observed in the characteristics extended area of the northeastern Peri-Tethys basin. Lithological and geochemical features of A study of the Dzhengutay section in sediments that correspond to the PETM vary Dagestan, East Caucasus (Fig. 1), shows a notably in an E-W direction from Central Asia regular fluctuation of TOC content within the to Crimea and in a S-N direction from SB. At least four 15–20 cm bands of black Transcaucasia to the central Russian Platform. shale and a lighter, more calcareous layer can In the deeper southern part of basin (Crimea, be recognized (Fig. 2). While the bottom and Caucasus and central Asia) it is characterized top of each band is very prominent, the by calcareous and calcareous-clayey transition from the lower black to upper pale sedimentation, and a sapropel bed (SB) with layers within the band is relatively gradual. varying TOC content that accumulated during Such a regular alternating pattern in lithology PETM. In the northern shallower area of the suggests a regular variation in sedimentary basin, siliceous-clayey sedimentation pattern of the SB formation. In the Kheu dominates and the PETM is diffiuclt to section, central Caucasus (Fig. 1), three cycles recognise, with a turnover in siliceous can be recognized — the most pronounced microfossils being the main tool to identify the (lower) cycle is ~25 cm and two additional PETM (Oreshkina and Oberhänsli 2003). The cycles are ~10 cm in thickness. Two cycles in thickness of the SB ranges from a few the SB are also distinguished in the Kurpai decimeters to a few meters in different areas, section, Tadjikistan. and the lithology and concentrations of TOC and trace elements vary greatly (Gavrilov et al. The cyclic nature of the SB is best defined in 1997, 2003; Gavrilov and Shcherbinina 2004). sections that contain a high TOC content (e.g. A recent study of new sections in different not less than 5%). In sections where TOC is parts of the basin that span the PETM (Fig. 1) lower (e.g. 2–3%), the variations in TOC found evidence of erosion below the SB that, content are of a lower magnitude and cyclicity once again, confirmed the occurrence of a is poorly defined. Only one cycle is present in regressive pulse prior to the PETM and that accumulation of the SB formed during a rapid transgression (Gavrilov et al. 1997). Moreover, the combination of results obtained from new sections and a revision of previously studied sections has provided more information regarding the SB Figure 1 Paleogeographic map showing localities of studied sections that contain architecture. the sapropel bed (SB) related to PETM. Blue – Peri-Tethyan basin, yellow – land, dashed area delineates a realm of occurrence of oil-shale coeval to SB. Extended Abstracts

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sections with low TOC content. Fluctuations of TOC within cycles are also accompanied by variations in many trace element concentrations (e.g. V, Ni, Mo, Se, Zn, Cu, U, Au, Ag). Spatial changes in the SB indicate that the concentration and nature of organic matter differs throughout the basin. The eastern part of the NE Peri-Tethys (Central Asia), an area with an arid climate, has the highest TOC concentrations (up to 25% and more) within the SB. Results from pyrolysis suggest the nature of organic matter is primarily from the marine basin, with a minor terrestrial component. In the western sector of the basin (Crimea and Caucasus), the TOC content is significantly lower (varies from 50% СаСО3) and the TOC content does not exceed 6%. Significant δ13С variations occur within the SB that can be linked with variations in TOC content and different parts of the cycle within the SB. For example, in the Kheu section (central Caucasus), the peak of the negative δ13С excursion (-1.8 ‰) occurs within the lower black shale (Fig. 2), which contains the highest TOC content (up to 10 %), while in the upper (pale) part of the same cycle the TOC drops up to ~1.0% and the magnitude of the δ13С excursion decreases (-0.6‰). The same trend is detected in the upper cycle. A similar pattern of δ13С fluctuations is also recognised in different cycles of the SB in the Dzhengutay section. The rate of onset and recovery of the δ13С excursion is different between sites. In some sections, a small decrease in δ13С occurs in sediments underlying the SB, but the most notable negative excursion occurs at the base of SB. Thus, in most sections, the onset of the δ13С excursion correlates to the base of the SB. The relationship between the top of the SB and δ13С recovery differs. The upper boundary of the SB is usually gradual, but nevertheless clearly pronounced, because the TOC content decreases rapidly. The negative δ13С excursion, however, continues for several decimeters or even meters above the horizon where TOC content drops. Hence, the δ13С excursion extends beyond the thickness of the SB. The δ18О excursion in the Crimean and Caucasian sections varies from 1.5–3‰, and in the Kheu section it reaches 5‰ (Fig. 2). In Central Asia, the magnitude of the excursion varies from 3.5‰ (Torangly, Turkmenistan) to 5‰ (Gur-Fatima, Tadjikistan) and 7‰ (Kurpai, Tadjikistan), while it is 2.5‰ in the Aktumsuk section. As with the δ13С record, in some sections the δ18О excursion also continues beyond the stratigraphic thickness of the SB. The magnitude of the δ13С and δ18О excursion differs throughout the basin — maximum values are recorded in the Central Asian sections, where sediments are enriched in organic matter and have low CaCO3 content. The Crimea and Caucasus areas were

Extended Abstracts


Figure 2 Lithology and contents of CaCO3, total organic carbon (TOC), and hydrogen index obtained from the results of organic matter (OM) pyrolysis (Rock-Eval II), carbon isotope composition from carbonate and OM, and oxygen isotope composition of Kheu (central Caucasus) and Dzhengutay (eastern Caucasus) sections.

characterized by a warm humid climate, whereas Central Asia had a hot arid climate. This may have affected the temperature of the shallow epeiric basin and resulted in the large δ18О excursion. At the same time, a high enrichment of organic matter in sediments may have enabled a high intensity of diagenetic processes that contributed to the large δ13С and δ18О excursions, particularly in calcareouspoor sediments. Nannofossil assemblage changes Nannofossil assemblage charactersistics of the SB, and recognised in different parts of the basin, include a dramatic decrease in the total nannofossil abundance, a wide occurrence of the short-lived species Discoster anartios, D. mahmoudii and rhomboasters (so-called “excursion taxa”), the elimination of coolwater chiasmoliths, an increase in warm-water discoasters and high-fertility Toweius, along with a decrease in Coccolithus spp.. Fasciculithus displays irregular behaviour in different areas. In some sections they distinctly increase in abundance (Torangly, Kurpai, Maly Zelenchuk, Dzhengutay), while in other sections they slightly decrease. Extended Abstracts

Generally, nannofossil assemblages from the SB are more abundant and diverse in shallower areas (Nasypnoe, Maly Zelenchuk, Torangly, Aktumsuk), characterized by lower TOC content, and the first occurrence of “excursion taxa” are coeval with the base of the SB. In the deeper parts of basin (Kheu, Dzhengutay) and eastern areas (Kurpai, Guru-Fatima), where the SB is more enriched in organic matter, nannofossils are more stressed and show rhythmical fluctuations in abundance that are coherent with the lithological cyclicity. In these sections, the lower parts of the SB (enriched in organic matter) contain extremely poor nannofossil assemblages and the lowest occurrence of “excursion taxa” is recorded in the middle or even upper part of the SB. This results in a minor diachroneity of the base of nannofossil Zone NP10 throughout the basin. Nannofossils in the SB from the eastern part of the basin are sparse and often form oligotaxonic assemblages largely dominated by Braarudosphaera bigelowii (Guru-Fatima section), or Fasciculithus spp. (Kurpai section), or Fasciculithus and Toweius spp. (Torangly section). Nannofossil assemblages appear to recover when accumulation of the

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SB ceases, and this recovery is independent of the continuation of the δ13C and δ18O excursion above the SB layer. DISCUSSION AND CONCLUSIONS Negative δ13C and δ18O excursions are documented in all studied sections, where they correspond to accumulation of the SB and substantial changes in nannofossil assemblages. However, the magnitude of the δ13C and δ18O excursion differs considerably throughout the basin, with the largest excursions recognised in Central Asia. The cyclic nature of the δ13C excursion within the SB horizon appears to co-vary with organic matter content. The magnitude of the isotope excursions was probably influenced by early diagenesis processes. Studies of several new sections from the NE Peri-Tethys reveal the cyclic architecture of the SB corresponds to the PETM, with accumulation of the SB occurring during the PETM (Gavrilov et al. 1997, 2003; Gavrilov and Shcherbinina 2004). This implies a rapid flooding of wide shoreland marshes, which enhanced the influx of biophile elements into the basin and resulted in unusual blooms of dinoflagellates, algae and bacteriaplankton (a kind of “red-tide”) that led to increased accumulation of organic matter on the basin floor. Recent studies of the cyclic nature of the SB (1–4 cycles and more) concluded that the transgression may have been interrupted by periodic short-lived still-stands, or even regressions, which reduced nutrient supply and led to productivity decay and accumulation of sediments with low amounts of organic matter. A number of cycles within the SB appear to have been controlled by local geomorphological features of the coastal landscapes. A wide low-gradient coastal area is the most likely landscape for formation of the cyclic SB. Alternatively, a coast with a steep slope (e.g. cliff) may have lowered the nutrient

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supply and resulted in reduced productivity, lower TOC in sediments and poorly-defined sedimentary cycles. Periodic interruptions of the transgression, which led to the cyclic nature of the SB formation, were more likely controlled by short Milankovich cycles. Examination of the PETM in the NE PeriThetys demonstrates this area was influenced by a mix of global (biotic and abiotic) and regional trends. ACKNOWLEDGEMENTS RFBR Project no. 09-05-00872. REFERENCES Gavrilov, Yu.O.; Kodina, L.A.; Lubchenko, I.Yu.; Muzylev, N.G. 1997: The Late Paleocene Anoxic Event in Epicontinental Seas of PeriTethys and Formation of Sapropelite Unite: Sedimentology and Geochemistry. Lithology and Mineral Resourses 35: 427-450. Gavrilov, Yu.O.; Shcherbinina, E.A.; Oberhänsli, H. 2003: Paleocene/Eocene boundary events in the northeastern Peri-Tethys. In Wing, S.L.; Gingerich, P.D.; Schmitz, B.; Thomas, E. ed. Causes and consequences of Globally Warm Climates in the Early Paleogene. Geological Society of America, Special Paper 369. Pp. 147-168. Gavrilov, Yu.O.; Shcherbinina, E.A. 2004: Global Biosphere event at the Paleocene-Eocene boundary. In Gavrilov, Yu.O.; Khutorskoy M.D. ed. Modern Problems of Geology. Moscow, Nauka. Pp. 493-531 (in Russian). Oreshkina, T.V.; Oberhänsli, H. 2003. Diatom turnover in the early paleogene diatomite of the Sengiley section, Middle Povolzhie, Russia: a response to Initial Eocene Thermal Maximum? In Wing, S.L.; Gingerich, P.D.; Schmitz, B.; Thomas, E. ed. Causes and consequences of Globally Warm Climates in the Early Paleogene. Geological Society of America, Special Paper 369. Pp. 169-180.

Extended Abstracts


MULTIPLE EARLY EOCENE THERMAL MAXIMA AT LOW LATITUDE PLATFORM CARBONATES – THE EOCENE SUCCESSION OF THE GALALA MOUNTAINS, EGYPT S. Hoentzsch1, C. Scheibner1, A. M. Marzouk2, M. W. Rasser3 and H.-J. Kuss1 1

2

Dept. of Geosciences, University Bremen, PO Box 330440, Bremen, Germany: [email protected]; Geol. Dept., Fac. of Science, Tanta University, Tanta 31527, Egypt; 3 Staatliches Museum für Naturkunde, D70191 Stuttgart, Germany

SUMMARY During the Early Eocene a large number of biotic and oceanic perturbations are recorded in deep sea environments that are underexplored with respect to their repercussions on shallow-marine carbonate platforms. This study provides new data for the Eocene succession on the isolated Galala carbonate platform in Egypt. We focus on the isotopic and geochemical evolution, as well as the general microfacies trends on a slope section. A detailed biostratigraphy based on calcareous nannofossils has provided a suitable base for the classification of multiple Eocene pertubations in the carbon cycle. The detected short-term, post-PETM hyperthermal events (e.g. ETM–2, ETM–3) in the Bir Dakhl area, are poorly evidenced from the southern Tethyan margin. We show preliminary results for δ13C, total carbon and total organic carbon pertubations and their impact on the depositional system. INTRODUCTION AND GEOLOGICAL SETTING The Galala Mountains are located in the Eastern Desert of Egypt and range from Ain Sukhna near Suez 100 km to the SE (Fig. 1a). The mountain complex represents an excellent example of an isolated Cretaceous to Eocene carbonate platform at the southern margin of the Tethys. The evolution of the platform is closely connected to the tectonic activity of the NE-SW striking Wadi Araba Fault as a part of the Syrian Arc-Fold-Belt (Hussein and AbdAllah 2001). Three major tectono-sedimentary segments can be distinguished (Fig. 1b, Kuss et al. 2000). The Northern Galala-Wadi Araba High (NGWA) represents shallow-marine inner platform environments. Due to synsedimentary uplift and erosion of the Wadi Araba Anticline Extended Abstracts

since the Upper Cretaceous, major inner platform outcrops are not accessible. The connection between NGWA and Southern Galala Subbasin (SGS) is represented by a transitional slope zone. The Eocene deposits of the Galalas can be subdivided in two lithostratigraphic units. The Thebes Group (Hermina and Lindenberg 1989) represents a deep water facies and is characterized by alternating chalky marls, cherts and sandstones. In contrast, the Southern Galala Formation (Abdallah et al. 1970; Kuss and Leppig 1989) consists of platform-related, shallow-marine limestones, sandstones and conglomerates. Both units interfinger in the upper slope of the southern Galala Mountains (Wadi B, Bir Dakhl). Former studies in the Galala area have focused on the pre-Eocene facies architecture and biostratigraphy (e.g., Kuss 1986; Bandel and Kuss 1987; Kuss et al. 2000; Scheibner et al. 2003), as well as on the biotic and climatic turnover around the Paleocene Eocene Thermal Maximum (PETM; Scheibner et al. 2005, Scheibner and Speijer 2008, 2009). In this study we concentrate on the investigation of the Eocene succession with respect to the Early Eocene Climatic Optimum (EECO). We try to detect post-PETM carbon isotope excursions (CIE) at the platform, which are related to a global shift towards negative carbon isotopes and warmer temperatures (Zachos et al. 2001). So far, these Eocene thermal maxima (ETM–2, ETM–3) have been described for deep-marine settings but have not been recognized in shallow-marine environments (e.g., Lourens et al. 2005; Dutton et al. 2005; Röhl et al. 2005). MATERIAL AND METHODS Nine sections, located on a N-S transect between Ain Sukhna and Ras Gharib in the

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Figure 1a Simplified map of the study area including the Paleocene–Eocene Global Boundary Stratotype Section and Point (GSSP) of Dababiya; 1b – Detailed map of the working area, showing the location of all sections, topographic features as well as the major tectono-sedimentary segments of the carbonate platform (modified after Kuss et al. 2000).

eastern Galala Mountains are recorded in detail (Fig. 1). Sections Wadi Waida 1, 2 (W1, W2) represent the NGWA; sections Wadi Ashkar 8 (K8), Wadi B 1/5 (B1/5) and Bir Dakhl 5, 6 (D5, D6) cover the slope and sections; sections Bir Dakhl 2, 4 (D2, D4) and Wadi Tarfa 3/4 (T3/4) represent the SGS. This manuscript concentrates on section D5, which has been continued from former investigations (Scheibner et al. 2003, 2005). The PETM interval of the section was resampled to get a better isotopic resolution of the event layer. A further continuation of section D5 is in progress. 143 samples, which cover 53 m of the lower Eocene succession of section D5, are taken for detailed geochemical and thin section analysis. The sample spacing

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varies between 20 cm in well-exposed marls and >1 m in sandstones, depending on the general outcrop conditions and alteration. Thirteen lithofacies types are defined in all studied sections regarding the distribution of the main faunal groups (smaller and larger benthic foraminifera, planktic foraminifera, green and red algae), matrix (micritic, sparitic, dolomitic), peloids and quartz. However, in section D5 only nine lithofacies types are present. Variations in the composition of larger benthic foraminifera communities and microfacies variations are used for a general Paleobathymetrical classification. Bulk rock geochemical analysis (total carbon - TC and total organic carbon - TOC) as well as isotope measurements (δ13C, δ18O) are conducted for all samples of section D5. Due to the Extended Abstracts


diagenetic alteration, only carbon isotopes are discussed below. Biostratigraphy is based on calcareous nannoplankton (NP zonation after Martini 1971), while a shallow benthic biostratigraphy based on larger foraminifera (Serra-Kiel et al. 1998) is in preparation. PLATFORM CONFIGURATION AND LATERAL FACIES TRENDS The restricted platform interior is characterized by strongly dolomitized to peloidal limestones as well as cyclicly deposited birdseye limestones, reworked conglomerates and rare marls (section W2). A similar succession in Upper Egypt was interpreted as tidal flat facies (Keheila and El-Ayyat 1990). Typical deposits from the platform interior are peloidal wacketo packstones with small lenticular nummulitids, green algae and miliolids. Lowenergy, open marine environments are

characterized by red algal bindstones (section B1/5) and Orbitolites-green algal floatstones (section W2). Sheltered backshoal deposits are alveolinid-rich wacke- to rudstones (section W2). High-energy shoals at the upper slope margin show accumulated nummulitids and alveolinids with varying quartz content (e.g. section B1/5). In slope settings, the occurrence of marls, cherts as well as sandstones and quartz-rich limestones increase. The abundance of larger foraminifera decreases with increasing water depth (sections D5 and D6). The basinal succession (section T3/4) is characterized by rhythmical bedded chalky marls and cherts; quartz grains and larger foraminifera are absent. The studied section D5 is recorded from prePETM NP9 to NP12. The succession is characterized by alternating marls, quartz-rich limestones and debris flow related sandstones

Figure 2 The succession of section D5 including the main lithology, lithofacies types, TC, TOC, δ13C, components and the general fossil content. Note the continuous negative trend in δ13C from the upper Paleocene (NP9) to lower Eocene (NP11/12, black arrow). Lithofacies signature shifts from shallow-marine platform related deposits (e.g. peloidal pack- to grainstones) to deeper marine slope facies (hemipelagic marls and quartz-rich limestones). Extended Abstracts

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and conglomerates. The abundance of marls increases towards the top of the section.

probably indicates highly conditions prior to the ETM–3.

EARLY EOCENE THERMAL MAXIMA – A CASE STUDY

Both ETM–2 and ETM–3 occur in isolated marl beds within sand- and quartz-rich limestone intervals. Subsequent to the ETM–3 an increase of chert layers is obvious within the marls and may reflect raised concentrations of dissolved silica (Muttoni and Kent 2007).

Biostratigraphic and isotopic results are shown together with microfacies data for the slope section D5 (Fig. 2). A prominent trend towards negative δ13C values starts with the PETM and continues up to the base of NP12. This trend is similar to the trend of the combined carbon isotope curve of Zachos et al. (2001). A decrease in TC concentration around the PETM, reflecting a rapid acidification of the oceans (Zachos et al. 2005), is not detected. A significant positive excursion in TOC, up to 0.45%, is present in the PETM interval, which is almost ten times higher than the general background TOC concentration (around 0.05%). Three further negative δ13C excursions occur above the PETM, which are accompanied by negative peaks in TC and significant positive excursions in TOC. They indicate pertubations similar to the PETM and are superimposed on the general trend towards negative δ13C values (Fig.2). The yellow interval indicates the ETM–2 (ELMO–H1 event) at the base of NP11 (Cramer et al. 2003). The CIE of the ETM–2 exhibits a shift of 1.5‰, similar to the CIE in other studies (e.g., Kroon et al. 2007). The TC concentration in the ETM–2 interval shifts from 10% to 5%, while no significant change in TOC is observed. A second carbon isotope excursion in NP11 probably shows a subordinate thermal maximum related to the I1 event (Lourens et al. 2005, Nicolo et al. 2007). The ETM–3 (“x” event; Röhl et al. 2005) is present at the base of NP12 and marks the onset of the EECO (Zachos et al. 2001). In contrast to previous studies, the CIE of ETM–3 is more pronounced in the Bir Dakhl area (around 2‰) than in deep sea records (0.6– 1‰; Röhl et al. 2005, Kroon et al. 2007). TC and TOC concentrations show prominent excursions: the bulk rock TC concentration is depleted to 1% in the event layer while the TOC concentration increases to more than 0.5%. However, the major TOC peak in the ETM–3 interval does not coincide with the negative excursions in δ13C and TC, which

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condensed

CONCLUSIONS 1) The Eocene succession of the Galala Mountains is probably the first record of post-PETM hyperthermals in a shallowmarine platform setting. ETM–2 and ETM–3 are determined in section D5 within intervals of high siliciclastic input. The negative CIEs at ETM–2 and ETM–3 coincide with strongly depleted TC concentrations. Positive TOC excursions are present within the PETM and ETM–3 interval. To support the present dataset, high resolution studies on the relevent intervals are in progress. 2) Nine sections at the Galala platform were studied in detail. Thirteen lithofacies types and three lithofacies subtypes have been determined. Detailed biostratigraphy for deeper marine slope and basinal sections is based on calcareous nannofossils, while biostratigraphy for platform environments with larger benthic foraminifera is in progress. The Eocene succession at the Galala Mountains is dominated by larger benthic foraminifera (alveolinids, nummulitids, discocyclinids and assilinids) at the platform and planktonic foraminifera in the basinal intervals. Green algae, gastropods, miliolids and coralline red algae dominantly occur at the platform interior. REFERENCES Abdallah, A.M.; Sharkawi, M.A.E.; Marzouk, A. 1970: Geology of Mersa Thelmet area, southern Galala, Eastern Desert, A.R.E. Bull. Fac. Sci., Cairo Univ. 44: 271-280. Bandel, K.; Kuss, J. 1987: Depositional environment of the pre-rift sediments - Galala Heights (Gulf of Suez, Egypt). Berliner geowissenschaftliche Abhandlungen A 78: 148.

Extended Abstracts


Cramer, B.S.; Wright, J.D.; Kent, D.V.; Aubry, M.P. 2003: Orbital climate forcing of δ13C excursions in the late Paleocene-early Eocene (chrons C24n-C25n). Paleoceanography 18: NO. 4, 1097, doi:10.1029/2003PA000909, 2003. Dutton, A.; Lohmann, K.C.; Leckie, R.M. 2005: Insights from the Paleogene tropical Pacific: Foraminiferal stable isotope and element results from Site 1209, Shatsky Rise. Paleoceanography 20: PA3004, doi:101029/2004PA001098. Hermina, M.; Lindenberg, H.G. 1989: The Tertiary. In Hermina, M.; Klitzsch, E.; List, F.K. ed. Stratigraphic lexicon and explanatory notes to the geological map of Egypt 1:500 000. Cairo, Conoco Inc. Pp. 141-217. Hussein, I.M.; Abd-Allah, A.M.A. 2001: Tectonic evolution of the northeastern part of the African continental margin, Egypt. Journal of African Earth Sciences 33: 49-68. Keheila, E.A.; El-Ayyat, A.A.M. 1990: Lower Eocene carbonate facies, environments and sedimentary cycles in Upper Egypt: evidence for global sea-level changes. Palaeogeography, Palaeoclimatology, Palaeoecology 81: 333-47. Kroon, D.; Zachos, J.C.; Leg 208 Scientific Party 2007: Leg 208 synthesis: Cenozoic climate cycles and excursions. In Kroon, D.; Zachos, J C.; Richter, C. ed. Proc. ODP, Sci. Results, 208. Pp. 1-55, doi:10.2973/odp.proc.sr.208.201.2007. Kuss, J. 1986: Facies development of Upper Cretaceous-Lower Tertiary sediments from the monastery of St. Anthony/Eastern Desert, Egypt. Facies 15: 177-194. Kuss, J.; Leppig, U. 1989: The early Tertiary (middle-late Paleocene) limestones from the western Gulf of Suez, Egypt. Neues Jahrbuch für Geologie und Paläontologie, Abhandlungen 177 (3): 289-332. Kuss, J.; Scheibner, C.; Gietl, R. 2000: Carbonate platform to basin transition along an Upper Cretaceous to Lower Tertiary Syrian Arc Uplift, Galala Plateaus, Eastern Desert, Egypt. GeoArabia 5: 405-424. Lourens, L.J.; Sluijs, A.; Kroon, D.; Zachos, J.C.; Thomas, E.; Röhl, U.; Bowles, J.; Raffi, I. 2005: Astronomical pacing of late Palaeocene to early Eocene global warming events. Nature 435: 1083-1087. Martini, E. 1971: Standard Tertiary and Quaternary calcareous nannoplankton zonation. In Farinacci, A. ed. Proceedings of the II Plankton

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Conference, Roma. Roma, Tecnoscienza Rome. Pp. 739-785.

Edizioni

Muttoni, G.; Kent, D.V. 2007: Widespread formation of cherts during the early Eocene climate optimum. Palaegeography, Palaeoclimatology, Palaeoecology 253: 348362. Nicolo, M.J.; Dickens, G.R.; Hollis, C.J.; Zachos, J.C. 2007: Multiple Early Eocene hyperthermals: Their sedimentary expression on the New Zealand continental margin and in the deep sea. Geology 35: 699-702. Röhl, U.; Westerhold, T.; Monechi, S.; Thomas, E.; Zachos, J.C.; Donner, B. 2005: The third and final Early Eocene thermal maximum: characteristics, timing, and mechanisms of the "X"-event. GSA Abstracts with Programs 37: 264. Scheibner, C.; Reijmer, J.J.G.; Marzouk, A.M.; Speijer, R.P.; Kuss, J. 2003: From platform to basin: The evolution of a Paleocene carbonate margin (Eastern Desert, Egypt). International Journal of Earth Sciences 92: 624-640. Scheibner, C.; Speijer, R.P.; Marzouk, A.M. 2005: Larger foraminiferal turnover during the Paleocene/Eocene thermal maximum and paleoclimatic control on the evolution of platform ecosystems. Geology 33: 493-496. Scheibner, C.; Speijer, R.P. 2008: Late Paleocene– early Eocene Tethyan carbonate platform evolution — A response to long- and shortterm paleoclimatic change. Earth Science Reviews 90: 71-102. Scheibner, C.; Speijer, R.P. 2009: Recalibration of the Tethyan shallow-benthic zonation across the Paleocene-Eocene boundary; the Egyptian record. Geologica Acta 7: 195-214. Serra-Kiel, J.; Hottinger, L.; Caus, E.; Drobne, K.; Ferrandez, C.; Jauhri, A.K.; Less, G.; Pavlovec, R.; Pignatti, J.; Samso, J.M.; Schaub, H.; Sirel, E.; Strougo, A.; Tambareau, Y.; Tosquella, J.; Zakrevskaya, E. 1998: Larger foraminiferal biostratigraphy of the Tethyan Paleocene and Eocene. Bulletin de la Société Géologique de France 169: 281-299. Zachos, J.; Pagani, M.; Sloan, L.; Thomas, E.; Billups, K. 2001: Trends, rhythms, and aberrations in global climate 65 Ma to present. Science 292: 686-693. Zachos, J.C.; Röhl, U.; Schellenberg, S.A.; Sluijs, A.; Hodell, D.A.; Kelly, D.C.; Thomas, E.; Nicolo, M.; I Raffi; Lourens, L.J.; McCarren, H.; Kroon, D. 2005: Rapid acidification of the ocean during the Paleocene-Eocene Thermal Maximum. Science 308: 1611-1615.

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CLIMATE SEE-SAWS IN PALEOGENE NEW ZEALAND Christopher J. Hollis1, Matthew Huber2, Luke Handley3, Kyle Taylor3, Richard D. Pancost3, John Creech4, Joel Baker4, Stefan Schouten5, Erica Crouch1 and Poul Schioler1 1

GNS Science, Lower Hutt, New Zealand: [email protected]; 2Earth and Atmospheric Sciences Department, Purdue University, USA; 3Bristol Biogeochemistry Research Centre, University of Bristol, UK; 4School of Geography, Environment and Earth Sciences, Victoria University of Wellington, New Zealand; 5Dept of Marine Biogeochemistry and Toxicology, Royal Netherlands Institute for Sea Research, The Netherlands.

New multi-proxy records of sea and air temperature variation from the upper Paleocene and Eocene bathyal marine sedimentary succession in Canterbury Basin, south-eastern New Zealand, and elsewhere in the southwest Pacific suggest a much more dramatic climate history for the region than the mild climatic changes previously inferred from local oxygen isotope records, which are now known to have been compromised by diagenesis. Sea temperature estimates derived from wellpreserved marine microfossils and sediments in northern Canterbury indicate that both sea floor and sea surface temperatures (SFTs and SSTs) increased by >10°C from Late Paleocene to Early Eocene times (Fig. 1). Late Paleocene TEX86-derived SSTs of ca 19°C and GDGT-derived Mean Annual Air Temperature (MAAT) estimates of 10–15°C are remarkably consistent with coeval TEX86 and GDGT records from the South Tasman Rise (Bijl et al., in press) and indicate a cool to warm temperate climate for the region at 59–56 Ma. This time period corresponds with the Paleocene carbon isotope maximum (PCIM), defined by a pronounced positive excursion in benthic δ13C, which appears to reflect an episode of increased carbon burial, either on land or in the oceans, and may signal a significant drawdown of atmospheric CO2. Moreover, at ODP Site 1121 on the margin of Campbell Plateau, we have determined SFTs of ca 4°C from Mg/Ca ratios obtained by laser ablation mass spectrometry on well-preserved benthic foraminifera. At this bathyal site, these deepwater temperatures are comparable with the modern day Deep Western Boundary Current, a flow that was thought to have begun in the earliest Oligocene at the initiation of circumpolar circulation and major Antarctic glaciation. We present evidence for widespread deepwater scour and non-deposition, biosiliceous sedimentation and eustatic sea level fall (Schiøler et al., in press) that points 76

to late Paleocene glaciation of Antarctica prior to the opening of the Tasmanian or Drake gateways. In dramatic contrast, our multi-proxy study (TEX86, GDGT, Mg/Ca, δ18O) has shown that the Early Eocene of northern Canterbury experienced truly tropical conditions (Hollis et al., 2009). Further detailed Mg/Ca analyses and utilisation of a new TEX86 calibration (Liu et al., 2009) has improved the agreement between proxies. SST peaked at 30°C at ca 50 Ma and declined to 27°C by 48 Ma. Similarly, SFT peaked at 20°C at 50 Ma and declined to 14°C by 48 Ma. A small cooling step at 48.5 Ma appears to correspond to intensification in corrosive bottom water flow over the Campbell Plateau, as evident from the disappearance of the planktic-benthic δ18O offset at DSDP 277. The SST estimates are consistent with TEX86derived estimates from the South Tasman Rise (Bijl et al., in press) and are also in agreement with latest Eocene estimates (TEX86 and UK’37) for site 277 (Liu et al., 2009). Modelled ocean circulation patterns and sea temperatures under Early Paleogene greenhouse conditions (2240 ppmv CO2, NCAR CCSM3) suggest temperate conditions for New Zealand (mean annual SSTs of 1520°C), which is consistent with Paleocene temperature estimates but not with Eocene estimates. A tropical climate for Early Eocene New Zealand implies that the role of ocean currents or other mechanisms of poleward heat transport or high-latitude heat retention are much greater under hyper-greenhouse conditions than is allowed for in the current generation of climate models. REFERENCES Bijl, P.; Schouten, S.; Brinkhuis, H.; Sluijs, A.; Reichart, G.-J.; Zachos, J.C. 2009: Palaeogene temperature evolution in the southwest Pacific: cooling the greenhouse. Nature 461: 776-779.

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Hollis, C.J.; Handley, L.; Crouch, E.M.; Morgans, H.E.G.; Baker, J.A.; Creech, J.; Collins, K.S.; Gibbs, S.J.; Huber, M.; Schouten, S.; Zachos, J.C.; Pancost, R.D. 2009. Tropical sea temperatures in the high-latitude South Pacific. Geology 37: 99-102. Liu, Z.; Pagani, M.; Zinniker, D.; DeConto, R.; Huber, M.; Brinkhuis, H.; Shah, S.R.; Leckie, R.M.; Pearson, A. 2009. Global Cooling During the Eocene-Oligocene Climate Transition. Science 323: 1187-1190.

Schiøler, P.; Rogers, K.; Sykes, R.; Hollis, C.J.; Ilg, B.; Meadows, D.; Roncaglia, L.; Uruski, C. in press. Palynofacies, organic geochemistry and depositional environment of the Tartan Formation (Late Paleocene), a potential source rock in the Great South Basin, New Zealand. Marine and Petroleum Geology. Zachos, J.C.; Pagani, M.; Sloan, L.C.; Thomas, E.; Billups, K. 2001. Trends, rhythms, and aberrations in global climate 65 Ma to present. Science 292: 686-693.

Figure 1 New sea temperature estimates based on TEX86, Mg/Ca ratios and δ18O records from bathyal marine sections in onshore Canterbury Basin (mid Waipara River [MW], Hampden Beach [HB]) and the eastern margin of the Campbell Plateau (ODP Site 1121) from Late Paleocene to Middle Eocene. For reference, δ18O records from DSDP Site 277, the global benthic foraminiferal compilation of Zachos et al. (2001), and qualitative estimate of climatic trends over the same time period based on paleontological data are also shown.

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THE WOOLWICH FORMATION OF CROYDON, S. LONDON, UK: A PETM FAUNA AND FLORA REDISCOVERED J.J. Hooker1, M.E. Collinson2, A.G. Lawson3, S. Tracey4 and J.A. Skipper5 1

Department of Palaeontology, Natural History Museum, London, SW7 5BD, UK: [email protected]; 2Earth Sciences Department, Royal Holloway University of London, Egham, TW20 0EX, UK; 329 Hale Drive, Mill Hill, London, NW7 3EL, UK; 4International Trust for Zoological Nomenclature, c/o Natural History Museum, London, SW7 5BD, UK; 5Geotechnical Consulting Group, 52A Cromwell Road, London SW7 5BS, UK.

INTRODUCTION In 1882, a cutting was dug for the then new Woodside and South Croydon railway, known as the Park Hill section. The cutting exposed in succession: Late Cretaceous Chalk with flints, Late Paleocene Thanet Formation and Upnor Formation (formerly Woolwich and Reading Bottom Bed) and Early Eocene Reading Formation, Woolwich Formation and Blackheath Beds (Klaassen 1883). The two basal units of the Woolwich Formation, a bluish clay with hard calcareous bands and a dark organic clay filling three small channels that incise the bluish clay, yielded terrestrially derived biota associated with a largely brackish indigenous mollusc fauna. The flora consisted mainly of fragmentary leaves with preserved cuticle in the bluish clay and in the northern channel fill. Fragmentary tree trunks were also recorded in the bluish clay. Of the fauna, particularly interesting were the bones of the giant flightless bird Gastornis klaasseni Newton, 1886, and an ulna of the pantodont mammal Coryphodon (Newton 1883), which were found in the central channel fill. It is evident that no screenwashing of sediment was performed and a letter from J.S. Gardner in Klaassen (1883: 238) regrets the absence of fruits and seeds, which would have allowed precise taxonomic determinations of the flora. Likewise, the only tetrapods found belong to animals weighing more than 100 kg. NEW EXCAVATIONS IN THE CENTRAL CHANNEL In 1998, the welcome opportunity came to reexcavate the northern end of the cutting (Sandilands cutting: National Grid Reference TQ339655), where the central channel was located. This was because the site was being redeveloped for the construction of the Croydon Tramlink. In a large rotational slip on the west wall of the cutting, most of the Woolwich Formation succession capped by 78

Blackheath Beds was exposed using a mechanical excavator (Fig. 1), but the base of the bluish clay was not reached. The central channel fill was bulk sampled (ca 1 tonne from the base, ca 500 kg from higher parts) and later screenwashed though a 0.5 mm sieve. The aim of retrieving a comprehensive biota was thus realised (see list below). The plants, mammals and other tetrapods will be treated in detail elsewhere.

Figure 1 Lithic log of 1998 excavation of Woolwich Formation and basal Blackheath Beds at the site of the central channel, Sandilands, Park Hill Cutting, Croydon. Scale divisions = 20 cm. * marks the position of the concentration of vertebrates and plants. A narrow inaccessible trench excavated below the logged section showed a further 85 cm of the basal unit.

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Mollusca, Gastropoda

•

‘Lactarius’ curvidorsalis (Frost, 1931) (Lactariidae), otoliths

•

‘Scombrops’ sinuosus (Pomatomidae), otoliths

Stinton,

1965 1900)

•

Clithon pisiforme (Férussac, 1823)

•

Brotia melanioides (J. Sowerby, 1816)

•

Tympanotonos funatus intermedius (J. Sowerby, 1816)

•

•

Diaphyodus sauvagei (Leriche, (Labridae), palates, teeth

Hydrobia websteri (Morris, 1854)

•

•

?Sphyraenodus sp. (Scombridae), tooth

Bythinella parkinsoni (Morris, 1854)

•

Anomalorbina sp.

•

Lymnaea sp.

•

Physa sp.

Mammalia •

Neoplagiaulacidae (Multituberculata)

•

Peradectes louisi Crochet, (Marsupialia: Peradectidae)

Mollusca, Bivalvia

•

?Leptictidae (Leptictida)

•

Mytilus mitchelli (Morris, 1854)

•

Pantolestidae indet. (Pantolesta)

•

Ostrea sp.

•

Ailuravinae (Rodentia: ‘Paramyidae’)

•

Polymesoda 1860)

•

Neomatronella luciannae Russell, Louis and Savage, 1975 (Amphilemuridae)

•

Scrobiculabra condamini (Morris, 1854)

•

?Amphilemuridae

•

Martesia (s.l.) sp.

•

Leptacodon sp. (Archonta: Nyctitheriidae)

•

Teredo (s.l.) sp.

•

Platychoerops Plesiadapidae)

Elasmobranchii (sharks and rays)

•

Melaneremia sp. (Primates: Omomyidae)

•

Sylvestrilamia teretidens (White, 1931), teeth

•

Palaeonictis gigantea de Blainville, 1842 (Creodonta: Oxyaenidae)

•

Hypolophodon sp., teeth, dermal denticles

dulwichiensis

(Rickman,

sp.

1979

(Plesiadapiformes:

Other tetrapods Actinopterygii (ray-finned bony fishes) •

Lepisosteus suessionensis Gervais, 1852 (Lepisosteidae), teeth, scales, vertebrae

•

Amia sp. (Amiidae), teeth

•

?Pterothrissus sp. (Pterothrissidae), otolith

•

Albula eppsi White and Frost, 1931 (Albulidae), otoliths, teeth

•

Phyllodus toliapicus Agassiz, (Phyllodontidae), palates, teeth

•

Anguilla sp. (Anguillidae), otolith

•

‘Argentina’ abbatiae (Argentinidae), otoliths

•

?Arius sp. (Ariidae), otolith

•

Indet. (Gadidae), otolith

•

‘Serranidarum’ serranoides 1965) (Serranidae), otoliths

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Stinton,

1839

1965

(Stinton,

These belong to urodele and anuran amphibians, turtles (including trionychid carapace plates), a lizard (jaw), crocodilians (including blunt tooth types attributable to Diplocynodon) and birds. Plants There are three distinctive elements, a seed of the resembling Trichosanthes Cucurbitaceae, a seed resembling Decodon of the Lythraceae and a seed which appears to be a new genus of Theaceae. These three occur in other PETM floras but are not known from other stratigraphic intervals in the UK Paleogene (Collinson and Cleal 2001). Other taxa belong to genera with a much wider stratigraphic range and include a small-seeded Vitis species (Vitaceae), endocarps of Palaeosinomenium sp. (Menispermaceae),

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endocarps of tribe Iodeae (Icacinaceae) (two species) and Rhamnospermum bilobatum Chandler (Incertae Sedis). There are also small nutlets (possibly Betulaceae) and a few other undetermined fruits/seeds. This remains a very low diversity flora (in spite of the very large sample volume processed). DISCUSSION The presence of the molluscs Pitharella rickmani (underlying bluish clay only) and Polymesoda dulwichiensis indicates that the Woolwich Formation succession at Croydon belongs to the upper shelly clays unit (= Upper Shelly beds of Page and Skipper 2000), to which these species are restricted (ST pers. obs.). Indeed, the upper part of the Croydon Woolwich succession, oyster-rich clays succeeded by lithified corbiculid-rich shell beds, is very similar to that of the upper shelly clays at Peckham, where the full succession of interdigitating Woolwich and Reading Formations resting on Upnor Formation was recorded (Berry and Cooper 1977). The organic channel fills and the bluish clay unit were not, however, present there and, together, are likely to represent local channel fill at the base of the upper shelly clays transgression (Klaassen 1883, Fig. 1; Collinson et al. 2003, Fig. 3; sequence 4 of Skipper 1999). The Apectodinium acme that proxies the CIE and therefore the Paleocene–Eocene Thermal Maximum (PETM) (e.g. Sluijs et al. 2008) has recently been found at nearby Brixton, in both the lower and upper shelly clays of the Woolwich Formation (Collinson et al. 2009). The central channel at Croydon therefore dates from the later part of the PETM. The majority of the molluscs and fish are related to taxa that inhabit brackish or coastal marine waters today. The channel is therefore considered to have been a brackish environment into which the land-based biota was transported. The main importance of this site is the cooccurrence of land-based tetrapods and plants within the PETM in Europe. Preliminary ecological interpretation of the mammals is that they represent a mix of dietary and locomotor types, with a number that can be classed as arboreal and scansorial. This supports the evidence from the plants, which indicate open woodland with lianas of tropical aspect.

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ACKNOWLEDGEMENTS Pierre Schreve for screenwashing the sediment; Richard Jenkins of Amey McAlpine for access to the site and for providing a mechanical digger for the excavation; Martin Newcombe and Warwick Reynolds of English Nature for access to an environmentally sensitive area; and Andy Currant, Pierre Schreve, John Cooper, Julia Day, Andy Gale, Paul Jeffery and David Polly for help with sampling. REFERENCES Berry, F.; Cooper, J. 1977: A temporary exposure of the Paludina Band (Woolwich Beds) at Peckham, South London. Tertiary Research 1: 77-82. Collinson, M.E.; Cleal, C.J. 2001: The palaeobotany of the Palaeocene and Palaeocene-Eocene transitional strata in Great Britain. In Cleal, C.J.; Thomas, B.A., Batten, D.J.; Collinson, M.E. ed. Mesozoic and Tertiary Palaeobotany of Great Britain. Peterborough, Joint Nature Conservation Committee. Geological Conservation Review Series 22: 155-184. Collinson, M.E.; Hooker, J.J.; Gröcke, D.R. 2003: Cobham Lignite Bed and penecontemporaneous macrofloras of southern England: a record of vegetation and fire across the Paleocene–Eocene Thermal Maximum. Geological Society of America Special Papers 369: 333-349. Collinson, M.E.; Steart, D.C.; Harrington, G.J.; Hooker, J.J.; Scott, A.C.; Allen, L.O.; Glasspool, I.J.; Gibbons, S.J. 2009: Palynological evidence of vegetation dynamics in response to palaeoenvironmental change across the onset of the Paleocene-Eocene Thermal Maximum at Cobham, southern England. Grana 48: 38-66. Klaassen, H.M. 1883: On a section of the Lower London Tertiaries at Park Hill, Croydon. Proceedings of the Geologists’ Association 8: 226-248. Newton, E.T. 1883: Note on Coryphodon Remains from the Woolwich Beds of the Park Hill Section, Croydon. Proceedings of the Geologists’ Association 8: 250-254, pl.3. Newton, E.T. 1886: On the remains of a gigantic species of bird (Gastornis klaasseni, n. sp.) from the Lower Eocene beds near Croydon. Transactions of the Zoological Society of London 12: 143-160, pls 28-29.

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Page, D.; Skipper, J.A. 2000: Lithological Characteristics of the Lambeth Group. Ground Engineering 33: 38-39. Skipper, J. 1999: The stratigraphy of the Lambeth Group (Palaeocene) of south east England. PhD thesis, Imperial College, London.

Extended Abstracts

Sluijs, A.; Brinkhuis, H.; Crouch, E.M.; John, C.M.; Handley, L.; Munsterman, D.; Bohaty, S.M.; Zachos, J.C.; Reichart, G.-J.; Schouten, S.; Pancost, R.D.; Sinninghe Damsté, J.S.; Welters, N.L.D.; Lotter, A.F.; Dickens, J.R. 2008: Eustatic variations during the PaleoceneEocene greenhouse world. Paleoceanography 23: PA4216, doi:10.1029/2008PA001615.

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HOW GLOBAL WARMING AFFECTS TROPICAL RAINFORESTS? A PALEOGENE PERSPECTIVE Carlos Jaramillo Smithsonian Tropical Research Institute, Apartado 0843-03092, Balboa, Ancon, Panamá: [email protected]

The consequences of global warming on tropical vegetation are unknown. Today, most tropical rainforest lives at temperatures below 27.5°C (Fig. 1). Many have argued that tropical communities live near their climatic optimum (Stoskopf 1981), and that a slight increase in temperature could be deleterious to them (Tewksbury et al. 2008). Empirical examples in earth history might help us understand the behaviour of tropical biota during past climate change. Past tropical temperatures during the Paleogene for a fossil forest in northern Colombia named Cerrejón have been estimated using leaf margin analysis (Herrera et al. 2005) and snake morphology (Head et al. 2009). Leaf

margin analysis, however, can give only a minimum estimate of paleotemperature for tropical forests, because the regression models used in the method lack a modern analogue for forest at temperatures above 28°C. Snake morphology suggests a temperature of 32°C for the middle Paleocene of Colombia (Head et al. 2009), which seems to agree with other proxies as well as with global circulation models for the Paleogene (Huber 2008). The fossil record of the tropics shows overall that tropical biotas were able to cope with high temperatures over extensive periods of time (several millions of years). Paleocene tropical forests from northern Colombia were similar in composition to modern tropical forests (Wing

Figure 1 Mean annual precipitation and mean annual temperature for 99 tropical rainforest sites across the tropics. Note how all of the tropical rainforest are below 28°C today. The fossil site from the Paleocene of Colombia, Cerrejón, does not have a modern analogue in mean annual temperature. Cerrejón site shows the error bars for the estimation of temperature (Head et al. 2009) and precipitation (Herrera et al. 2008b).

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et al. 2004). Most of the plant families that are abundant in the neotropical rainforest today were also abundant in the Paleocene, including legumes, Malvaceae, palms, Araceae, and Menispermaceae (Doria et al. 2008; Gomez et al. 2009; Herrera et al. 2008a; Jaramillo et al. 2007; Ramirez 2009). The forest also supported a rich fauna that included mammals, giant snakes, crocodiles, and giant turtles (Bloch et al. 2008; Cadena and Jaramillo 2006; Hasting et al. 2009; Head et al. 2009). Moreover, a subsequent warming, the longterm Eocene thermal maximum with temperatures in the tropics reaching 36– 37°C, correlates with an increase in tropical plant diversity (Jaramillo et al. 2006). In contrast, modern experimental studies have shown than plants suffer several deleterious effects under maximum daily temperatures associated with mean annual temperatures of 32–33°C. Some of those effects include a decrease in the rate of photosynthesis, a decrease in net production, an increased risk of photoinjury, and an increase in isoprene emissions (e.g. Lerdau and Throop 1999; Stoskopf, 1981). How, then, to explain the Cerrejón forest thriving at 32°C? The solution could rely on a combination of high CO2 and elevated precipitation. The Cerrejón fossil forest lived under high precipitation regimes — about 3.2 m of rain a year (Herrera et al. 2005, 2008b) — and CO2 levels much higher than those of today (Royer 2006). Experiments in greenhouses have shown that plants deal better with high temperatures under high levels of CO2 and precipitation (Aber et al. 2001; Berry and Bjorkman 1980; Niu et al. 2008). Perhaps, then, rainforest lineages already have the genetic variability to cope with elevated temperatures if they are living in high levels of CO2 and high rates of rainfall, as during the Paleogene. ACKNOWLEDGMENTS This project was supported by the Smithsonian Paleobiology Endowment Fund, the Fondo para la Investigacion de Ciencia y Tecnología Banco de la Republica de Colombia, the Unrestricted Endowments Smithsonian Institution Grants, National Science Foundation, and Carbones del Cerrejón LLC.

Extended Abstracts

Special thanks go to M. I. Barreto for her continuous support, and ideas. REFERENCES Aber, J.; Neilson, R.; McNulty, S.; Lenihan, J.M.; Bachelet, D.; Draper, R.J. 2001: Forest processes and global environmental change: Predicting the effects of individual and multiple stressors. BioScience 51: 735-751. Berry, J.; Bjorkman, O. 1980: Photosynthetic response and adaptation to temperature in higher plants. Annual Review of Plant Physiology 31: 491-543. Bloch, J.; Cadena, E.; Hasting, A.; Rincón, A.; Jaramillo, C. 2008: New vertebrate faunas from the Paleocene sediments of Colombia, northwestern South America. In Society of Vertebrate Paleontologists Annual Meeting: Cleveland, USA. The Society of Vertebrate Paleontology. Cadena, E.; Jaramillo, C. 2006: New Podocnemididae fossil turtles from the late Paleocene Cerrejon formation, Guajira Peninsula, Colombia. In 66th Annual Meeting of the Society of Vertebrate Paleontology, Ottawa, Canada. Doria, G.; Jaramillo, C.; Herrera, F. 2008: Menispermaceae from the Cerrejon formation, middle to late Paleocene. American Journal of Botany 95: 954-973. Gomez, N.; Jaramillo, C.; Herrera, F.; Wing, S.L.; Callejas, R. 2009: Palms (Arecaceae) from a Paleocene rainforest of northern Colombia: American Journal of Botany: in press. Hasting, A.; Bloch, J.; Cadena, E.; Jaramillo, C. 2009: A new small short-snouted Dyrosaurid (Crocodylomorpha, Mesoeucrocodylia) from the Paleocene of northeastern Colombia. Journal of Vertebrate Paleontology: in press. Head, J.; Bloch, J.; Hasting, A.; Bourque, J.; Cadena, E.; Herrera, F.; Polly, P.D.; Jaramillo, C. 2009: Giant boine snake from a Paleocene Neotropical rainforest indicates hotter past equatorial temperatures. Nature 457: 715-718. Herrera, F.; Wing, S.; Jaramillo, C.A. 2005: Warm (not hot) Tropics during the Late Paleocene: First Continental Evidence. In AGU, E.T. ed. AGU Annual Meeting, Volume 86, Fall Meet. Suppl., Abstract PP51C-0608. Herrera, F.; Jaramillo, C.; Dilcher, D.; Wing, S.L.; Gomez, C. 2008a: Fossil Araceae from a Paleocene Neotropical rainforest in Colombia. American Journal of Botany 95: 1-16.

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Herrera, F.; Wing, S.L.; Jaramillo, C.; Gomez, C.; Dilcher, D.; Wilf, P.; Labandeira, C.C. 2008b: Fossils from Colombia reveal tropical rainforest in the Paleocene greenhouse world. Terra Nostra 2008-2: 113. Huber, M. 2008: A Hotter Greenhouse? Science 321: 353-354. Jaramillo, C.; Pardo-Trujillo, A.; Rueda, M.; Harrington, G.; Bayona, G.; Torres, V.; Mora, G. 2007: Palynology of the Upper Paleocene Cerrejon Formation, northern Colombia. Palynology 31: 153-189. Jaramillo, C.; Rueda, M.; Mora, G. 2006: Cenozoic plant diversity in the Neotropics. Science 311: 1893-1896. Lerdau, M., T.; Throop, H., L. 1999: Isoprene emissions and photosynthesis in a tropical forest canopy: Implications for model development. Ecological Applications 109: 1109- 1117. Niu, S.; Wu, M.; Han, Y.; Xia, J.; Li, L.; Wan, S. 2008: Water-mediated responses of ecosystem

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carbon fluxes to climatic change in a temperate steppe. New Phytologist 177: 209–219. Ramirez, M. 2009: Afinidades taxonómicas de una Malvaceae del Paleoceno Superior (58–60 Ma) en el norte de Colombia [Undergraduate Senior Thesis]. Medellin, Universidad de Antioquia. Royer, D.L. 2006: CO2-forced climate thresholds during the Phanerozoic. Geochimica et Cosmochimica Acta 70: 5665–5675. Stoskopf, N. 1981: Understanding crop production: Upper Saddle River, New Jersey. PrenticeHall. 433 p. Tewksbury, J.J.; Huey, R.B.; Deutsch, C.A. 2008: Putting the heat on tropical animals: Science 320: Wing, S.; Herrera, F.; Jaramillo, C. 2004: A Paleocene flora from the Cerrejon formation, Guajira Peninsula, northeastern Colombia. In International Organization of Paleobotany, Seventh Quafrennial Conference, Argentina, Volume Abstracts: 146-147.

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EOCENE FLORA AND PALEOCLIMATE OF HAINAN ISLAND, SOUTH CHINA JianHua Jin1, T. M. Kodrul 2, A.B. Herman2 and WenBo Liao1 1

School of Life Sciences, Sun Yat-sen University, Guangzhou 510275, P. R. China: [email protected]; 2 Geological Institute, Russian Academy of Sciences, Moscow, Russia: [email protected]

INTRODUCTION During last few decades, Eocene floras of the world have been the focus of numerous Paleobotanical studies. However, Eocene floras from tropical latitudes, particularly from South-east Asia, are still poorly known. Study of these floras, which existed during the warmest epoch of the Cenozoic, is important for better understanding of the origin and high diversity of modern tropical forests of Asia. This paper represents the results of our study of the Eocene flora from Hainan Island, which is suitable for Paleofloristic and Paleoclimatic analysis. This flora provides crucial data for reconstruction of low latitude terrestrial Paleoecosystems and vegetation. The Changchang flora comes from the Changchang Basin. It is located near Jiazi Town, Qiongshan County, in the northern part of Hainan Island, South China. Paleogene deposits of the Changchang Basin are subdivided into three formations (Zhou and Chen 1988; Lei et al. 1992) — Changtou Formation (Paleocene), Changchang Formation (Eocene) and Wayao Formation (Eocene). The plant-bearing Changchang Formation consists of lower and upper members. The lower member, 250–300 m, is represented by sediments of lacustrine origin — mudstone, siltstone, sandstone and conglomerate of different colours. The upper (coaliferous) member, about 190–230 m, consists of clastic terrigenous and coal-bearing deposits probably formed in a lacustrine and alluvial environment. These deposits are represented by dark grey mudstone, grayish black coaly shale, brownish grey oil-bearing shale, yellowish brown, grayish yellow, grayish white muddy siltstone and sandstone, and coal. The latter were accumulated in swamps and oxbow lakes. Numerous well-preserved plant fossils were collected mainly in coaly shales, grey mudstones and siltstones in the upper member of the Changchang Formation. These deposits contain gastropod and bivalve mollusk shells Extended Abstracts

as well as fish bones and scales. Spore and pollen assemblages are also known from the upper member (Zhang 1980; Lei et al. 1992; Yao et al. 2009). TAXONOMIC COMPOSITION OF THE CHANGCHANG FLORA Before our study of the Changchang flora, there was only a preliminary report on plant megafossils from the Changchang Formation (Guo 1979), in which 10 species belonging to 9 genera were identified and described. Recently, Jin (2009) studied two fossil fruits belonging to Paleocarya sp. (Juglandaceae) and Acer cf. A. miofranchetii Hu et Chaney (Aceraceae). Palynological studies in the Changchang Basin started in the mid 1960s. In the late 1980s and early 1990s, Lei et al. (1992) studied spore and pollen assemblages from the Changtou, Changchang and Wayao Formations of the Changchang Basin. Based on this palynological data, Lei et al. (1992) dated the upper (coaliferous) member of the Changchang Formation as late Early Eocene to early Late Eocene. Yao et al. (2009) described the palynoflora of Changchang Basin, and compared it with that of the Middle–Late Eocene of Hunchun City (Jilin Province, North China) to reconstruct a Paleoclimate based on the results of the palynological study. In recent years we have collected numerous plant fossils in the upper (coaliferous) member of the Changchang Formation in Hainan Island (19º38'03N, 110º27'04E). Impressions and compressions of leaves, fruits, rhizomes and roots, as well as petrified wood remains are most abundant in the grey mudstone underlying and overlying two main coal seams. The Changchang flora yields horsetails(?), ferns, conifers and angiosperms belonging to at least 170 species (morphotaxa). With respect to diversity, the Changchang flora is comparable to the richest Eocene floras of the world. Possible horsetails are represented by numerous rhizome remains probably belonging to Equisetum (Equisetaceae). Ferns are also not

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diverse in the flora. The most abundant fern is Osmunda lignitum which probably inhabited wetlands and swamps. Aquatic ferns are represented by rare Salvinia. Only a few remains of Dryopteridaceae-like ferns and one segment of fern probably belonging to Polypodiaceae were found. Conifers of the Changchang Flora are extremely rare. In the collection there are only a few compressions of podocarpaceous leaves (e.g. Nageia).

Eurasia during the Miocene (Wang et al. 2007). Family Malvaceae is represented by the genus Craigia (Jin et al. 2009). This is the first fossil record of Craigia in its modern distribution centre that provides a new data on the origin and phytogeographic history of this genus. Juglandaceous winged fruit was assigned to Palaeocarya sp. (Jin 2009). Family Ulmaceae is represented by leaves belonging to Celtis and possibly Ulmus.

Angiosperms, both dicots and monocots, dominate in the flora. Dicotyledons belong to the families Lauraceae, Nelumbonaceae, Fagaceae, Altingiaceae, Myricaceae, Fabaceae, Malvaceae, Juglandaceae and Ulmaceae.

Monocotyledons are the characteristic component of the Changchang Flora. Among them, palm leaves are the most numerous. These leaves were assigned by Guo (1979) to two species of Sabalites (Arecaceae). However, our recent find of fossil palm leaves suggest that they probably all belong to a single polymorphic species. Fragmentary elongate leaves with parallel venation possibly belong to Musaceae (Musophyllum) and Araceae.

Lauraceae are an almost ubiquitous component of the Changchang flora. Based on leaf morphology and epidermal-cuticular characters, several genera of this family could be recognised. In several localities, the most abundant plant remains are Nelumbo (Nelumbonaceae) leaves, rhizomes, roots, tubers, receptacle and fruits. New material, collected recently, has allowed us to reinterpret Cyclocarya scutellata, described by Guo (1979), as a Nelumbo receptacle. Family Fagaceae includes several species belonging to two or three genera (possibly Castanopsis, Lithocarpus and Quercus). Angiosperm family Altingiaceae is documented by three-lobed Liquidambar leaves similar to those of L. miosinica Hu et Chaney. Family Myricaceae is represented in the Changchang Flora by a single species of Myrica. Family Fabaceae is widely represented in the Changchang flora. Leaflets with asymmetrical base and short petiolule were preliminary assigned to morphogenus Leguminosites. Other leaflets possessing lanceolate lamina, rounded mucronate apex, asymmetric base and prominent basal veins, extending along leaflet lamina margins, possibly belong to the genus Podocarpium. Single seeded fruits associated with these leaflets are assigned to Podocarpium as well. This genus is known in the early Oligocene to Pliocene floras of Eurasia (Wang et al. 2007). New finds of Podocarpium in the Eocene of Hainan Island support the assumption that an ancestral population of this genus may have originated in the Early Paleogene of eastern Asia and then spread to most of middle latitude areas in

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Petrified wood is also common in the Changchang locality. At least two species of morphogenus Paraphyllanthoxylon similar in anatomy to the modern Elaeocarpaceae or Euphorbiaceae were recognized (Feng et al. submitted). A new species of Liquidambaroxylon has been described (Oskolsky et al. submitted). PALEOCLIMATIC INTERPRETATION OF THE CHANGCHANG FLORA Using the coexistence approach on the palynoflora of the Changchang Basin, Yao et al. (2009) estimated a mean annual temperature (MAT) at 14.2–19.8°C. However, leaf margin analysis of 147 non-aquatic dicot leaf morphotypes from the Changchang Flora shows that 76% have leaves with entire margins. This indicates that the estimated MAT experienced by the flora is approximately 24°C using the correlation of Wofe (1979). Such MAT allows us to interpret the Changchang vegetation as a paratropical rainforest. This is corroborated by abundance in the flora of lauraceous plants and by wood anatomy — growth rings of the Changchang woods are absent or indistinct (Oskolsky pers. comm.), which is a characteristic feature of the majority of plants living in hot (tropical) climate without a pronounced temperature or moisture seasonality. Humid conditions are

Extended Abstracts


corroborated by coal deposits in the Eocene of the Changchang Basin. ACKNOWLEDGEMENTS This study was supported by the National Natural Science Foundation of China (Grant No. 40672017), the NSFC-RFBR (Grants Nos. 40811120033 and 07-04-92127), the Key project of the Sun Yat-sen University for inviting foreign teachers, and the Guangdong Provincial Natural Science Foundation of China (Grant No. 06023161). We are grateful to graduate students of Sun Yat-sen University for collaboration in our field works in Hainan Island, to Alexei Oskolsky (Botanical Institute, Russian Academy of Sciences) for providing us with his data on Changchang fossil woods and to Christa Hofmann (University of Vienna) for her examination of some palynological samples. REFERENCES Feng, X.X.; Yi, T.M.; Jin, J.H. submitted: First record of Paraphyllanthoxylon fossil wood from China, IAWA Journal. Guo, S.X. 1979: Late Cretaceous and Early Tertiary floras from the southern Guangdong and Guangxi with their stratigraphic significance. In Mesozoic and Cenozoic red beds of South China. Beijing, Science Press. Pp. 223–230 (in Chinese). Jin, J. H. 2009: Two Eocene fossil fruits from the Changchang Basin of Hainan Island, China. Review of Palaeobotany and Palynology 153: 150–152. Jin, J. H.; Kodrul T.M.; Liao W.B.; Wang X. 2009: A new species of Craigia from the Eocene

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Changchang Formation of Hainan Island, China. Review of Palaeobotany and Palynology. ISS: 80-82 Lei, Y. Z.; Zhang, Q. R.; He, W.; Cao, X. P. 1992: Tertiary. In Geology of Hainan Island, I. Stratigraphy and palaeontology. Beijing, Geological Publishing House. Pp. 218-266 (in Chinese). Oskolsky, A.A.; Kodrul, T.M.; Jin, J.H., submitted: Leaves and wood of Altingiaceae from the Eocene Changchang Formation, Hainan Island (China). Paleontological Journal. Wang, Q; Dilcher, D.L.; Lott, T. 2007: Podocarpium A. Braun ex Stizenberger 1851 from the middle Miocene of Eastern China, and its palaeoecology and biogeography. Acta Palaeobotanica 47(1): 237–251. Wolfe, J. A. 1979: Temperature parameters of humid to mesic forests of Eastern Asia and relation to forests of other regions of the Northern Hemisphere and Australia. U. S. Geol. Survey Professional Paper 1106: 1–37. Yao, Y.F.; Bera, S.; Ferguson, D.K.; Mosbrugger, V.; Paudayal, K.N.; Jin, J.H.; Li, C.S. 2009: Reconstruction of paleovegetation and paleoclimate in the Early and Middle Eocene, Hainan Island, China. Climatic Change 92: 169–189. Zhang, Q.R. 1980: Stratigraphy and palaeontology. Yichang, Yichang Institute of Geology and Mineral Resources, Chinese Academy of Geological Sciences. Pp. 106–117 (in Chinese). Zhou, G.Q.; Chen, P.Q. 1988: Tertiary. In Regional Geology of Guangdong Province, People's Republic of China. Geological Publishing House, Beijing, Pp. 237–263 (in Chinese with English abstract).

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BARTONIAN–RUPELIAN STRATIGRAPHY IN SOUTHERN AUSTRALIA: THE LAST DAYS OF THE AUSTRALO-ANTARCTIC GULF Brian McGowran Earth and Environmental Sciences, The University of Adelaide, SA 5005 Australia: [email protected]

INTRODUCTION

CAESURA I

Australia-Antarctica separation produced the birth and widening of the Australo-Antarctic Gulf (AAG) and then its death when subsumed in the new Southern Ocean. Two views pertaining to the greenhouse-icehouse transition in the Paleogene are the “Dinocyst biogeographic hypothesis” (Huber et al. 2004) and the venerable “Tasman gateway hypothesis” (Exon and Kennett 2004). The critical stratigraphic interval, late Middle Eocene (latest Lutetian) to Early Oligocene (Rupelian), not coincidentally is also the main focus of the role of CO2 in the environmental transition. I have urged the appellation Auversian Facies Shift (Berger and Wefer 1996) for the interval, not only paleoceanographically as originally conceived but more broadly, capturing the sense of major change in geohistory and biohistory.

Late Lutetian (ca 42 Ma), the beginning of the last ca 10 m.y. for the AAG — rejuvenation globally of oceanfloor spreading, AAG widening regionally. In response were the Khirthar marine transgression and the onset of neritic carbonate accumulation in the Bartonian Age, simultaneously on the southern, western and northern Australian margins, the most significant transgression in the Indo-Pacific province, and probably triggering the onset of the Middle Eocene Climatic Optimum (MECO; Fig. 2). Photosymbiotic foraminifera and dinoflagellates identify two warming spikes on the proto-Leeuwin Current, one Wilson Bluff (coeval with MECO, and on terrestrialpaleobotanical evidence the warmest time in southern Australia), and the other Tortachilla. It was indeed warm and wet at ~60°S, with partial return of the deep chemical weathering conditions of the Early Eocene. Coevally with the neritic carbonates, world-class coal swamps developed, the biggest in the Latrobe Valley in the Gippsland Basin (Holdgate et al. 2009), and are also distributed along the AAG

There are two central themes here. The evidence for correlations and stratigraphic summarizing statements (Fig 1) is cited and reviewed in McGowran (2009) and McGowran et al. (2004).

Figure 1 The stratigraphic record in the oceanic, terrestrial, but especially neritic environmental realms on the north flank of the Australo-Antarctic Gulf (AAG) can be boiled down to six named marine transgressions, Wilson Bluff to Aldinga. Each transgression characterises one of six stratal packages that are acting as informal regional stages. They comprise a natural three-part succession, Bartonian–Priabonian–Rupelian, with three punctuating caesuras, I, II, III.

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Extended Abstracts


Figure 2 The regional AAG record is a microcosm in chronostratigraphic and successional respects of the global geo- and biohistorical succession. The six packages match “perfectly” (thereby strongly supporting) the putatively global third-order sequences Lu 4 to Ru 1 (Fig. 1). The three-part configuration likewise has a global feel, especially the Bartonian–Priabonian transition, hitherto somewhat under-emphasized.

margin. The gymnosperm-dominated rainforests of the Paleocene and Early Eocene had been altered by the advent of angiosperms, especially the “southern beech” genus Nothofagus.

Neritic microfaunas underwent pronounced biogeographic contraction and subsequent expansion and diversification (Moss and McGowran 2003). But coal swamps were still prominent.

CAESURA II

In scenarios for the death of the Eocene AAG, the Oligocene birth of the Southern Ocean and the great environmental transition, the neritic record of the northern AAG margin and evidence of an influential proto-Leeuwin Current are more in accord with the “Dinocyst biogeographic hypothesis”.

The Bartonian–Priabonian transition is marked by contraction and withdrawal of neritic carbonate facies and oligotrophic taxa, and expansion of grey-green-black sediments with characteristic infaunal dominances among dysaerobic and/or high-nutrient biotas. The “extreme” cases are the silicas—spiculites, spongolites, and opal-A/-CT alternations in marginal, barred or otherwise protected environments. Conditions were somewhat cooler than in the Bartonian and even wetter, with a widespread shift to estuarine circulation (i.e. surface-water outflow). More coals accumulated in the Latrobe Valley and to the west. Nothofagus further replaced the rainforest gymnosperms (Holdgate et al. 2009). CAESURA III The Eocene–Oligocene (=Priabonian– Rupelian) boundary and glaciation Oi-1 — regional downcuts of ~50m and backfills. Warm/wet conditions shifted toward cool/dry, with return to improved ventilation in neritic environments and resumption of carbonate accumulation in the neritic and oceanic realms. Extended Abstracts

DISCUSSION Caesuras I and III are physically more apparent and have received more attention, especially at the Eocene–Oligocene boundary, but it is becoming apparent that Caesura II is more pronounced in the fossil record. Within the Bartonian–Priabonian transition, also including Caesura II in the AAG, we find major biotic (chronofaunal) overturns in the neritic, pelagic and terrestrial realms. The Tethyan neritic shows a wholesale increase in available nutrients, Bartonian to Priabonian. In response, the large, photosymbiotic, benthic foraminifera suffered the most extensive change in their Paleogene history between shallow-benthic zones SBZ17 and SBZ18. The photosymbiotic planktonic foraminifera went extinct abruptly in the pelagic realm between zones E13 and E14, close in time at least to

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their ecological counterparts’ crisis in the neritic and probably simultaneously. In the deep-ocean benthics, the Priabonian turnover is heralded by a dysaerobic facies in Tethys. In the North American mammals, the most farreaching turnover was in the Duchesnean age. Comprehensive patterns for the Auversian Facies Shift demand comprehensive explanations embracing all three environmental realms. In the AAG the most obvious need is more quantified palynology, preferably both referring to and testing the six stratal packages and addressing the notion of really extensive shifts to estuarine circulation in the Priabonian. For example, the Priabonian estuarine circulation advocated in McGowran (2009) and in this extended abstract is strongly supported by dinoflagellate patterns in ODP holes just beyond the head of the AAG (Sluijs et al. 2003). The Priabonian Deflandrea and Phthanoperidinium assemblages were interpreted as reflecting nutrient supply dominated by runoff. Towards Caesura III, and just after the initial deepening of the Tasman Gateway, the assemblages were replaced by the Brigantedineum assemblage, interpreted as reflecting nutrient supply dominated by upwelling. Correlating and comparing this succession with the palynological record of neighbouring rainforests (Holdgate et al. 2009) will require rigorous but rewarding work. REFERENCES Berger, W.H.; Wefer, G. 1996: Expeditions into the past: paleoceanographic studies in the South Atlantic. In Wefer, G.: Berger, W.H.: Siedler, G: Webb, D.J. ed. The South Atlantic: present and past circulation. Berlin, Springer-Verlag. Pp. 363-410. Holdgate, G.R.; McGowran, B.; Fromhold, T.; Wagstaff, B.E.; Gallagher, S.J.; Wallace, M.W.; Sluiter, I.R.K.; Whitelaw, M. 2009: Eocene-Miocene carbon-isotope and floral

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record from brown coal seams in the Gippsland Basin of southeast Australia. Global and Planetary Change 65: 89-103, doi:10.1016/j.gloplacha.2008.11.001. Huber, M.; Brinkhuis, H.; Stickley, C.E.; Döös, K.; Sluijs, A.; Warnaar, J.; Schellenberg, S.A.; Williams, G.L. 2004: Eocene circulation of the Southern Ocean: was Antarctica kept warm by subtropical waters? Paleoceanography 19: PA4026, doi:10.1029/2004PA001014. Kennett, J.P.; Exon, N.F. 2004: Paleoceanographic evolution of the Tasman Seaway and its climatic implications. In Exon, N.F.; Kennett, J.P.; Malone, M.J. ed. The Cenozoic Southern Ocean: tectonics, sedimentation, and climate change between Australia and Antarctica. American Geophysical Union, Geophysical Monograph 151. Pp. 345-367. McGowran, B. 2009: The Australo-Antarctic Gulf and the Auversian Facies Shift. In Koeberl, C.; Montanari, A. ed. The Late Eocene Earth— Hothouse, Icehouse, and Impacts, Geological Society of America, Special Paper 452, doi: 10.1130/2009.2452(14). Pp. 215-240. McGowran, B.; Holdgate, G.R.; Li, Q.; Gallagher, S.J. 2004: Cenozoic stratigraphic succession in southeastern Australia. Australian Journal of Earth Sciences 51: 459-496. Moss, G.; McGowran, B. 2003: Oligocene neritic foraminifera in Southern Australia: spatiotemporal biotic patterns reflect sequencestratigraphic environmental patterns. In Olson, H.; Leckie, M. ed. Paleobiological, geochemical, and other proxies of sea level change. SEPM (Society of Sedimentary Geology) Special Volume 75: 117-138. Sluijs, A.; Brinkhuis, H.; Stickley, C.E.; Warnaar, J.; Williams, G.L.; Fuller, M. 2003: Dinoflagellate cysts from the Eocene/Oligocene transition in the Southern Ocean: results from ODP Leg 189. In Exon, N.F.; Kennett, J.P.; Malone, M.J. ed. Proceedings of the Ocean Drilling Program, Scientific Results 189: 1-42 (Online).

Extended Abstracts


STRATIGRAPHY OF TWO KEY PALEOGENE SECTIONS FROM CANTERBURY BASIN, NEW ZEALAND H.E.G. Morgans, C.J. Hollis and E.M. Crouch GNS Science, PO Box 30368, Lower Hutt, New Zealand: [email protected]

INTRODUCTION New Zealand is currently a focus of research for understanding paleoclimate during the Paleogene. In the Eocene, the New Zealand micro-continent occupied a position of ca 50– 65°S, where its sedimentary basins recorded the interaction between the warm north and the cool south. Canterbury Basin had a southerly aspect during the Eocene with tectonic reconstructions indicating it faced open ocean from the ESE to SSE (Fig. 1). Sea surface temperatures of 23– 25°C (Burgess et al. 2008) and 30°C (Hollis et al. 2009) from

two key Paleogene sections (Fig.2) in the Canterbury Basin are evidence for warm subtropical to tropical climate through the Early to Middle Eocene at paleolatitude ca 55°S in the New Zealand region. MID-WAIPARA RIVER The mid-Waipara River section, in the north of the Canterbury Basin, is a sequence of well exposed Late Cretaceous–Paleocene mudstones and siltstones, Late Paleocene greensands, Early to Middle Eocene bathyal mudstones, and Late Eocene–Oligocene

Figure 1 Paleogeographic reconstruction of the New Zealand micro-continent at 55 Ma (earliest Eocene), after Hollis et al. (2006). CI = Chatham Islands; HB = Hampden Beach; MS = Mead Stream; MW = mid-Waipara; EAC? = East Australia Current?; SAW = sub-Antarctic water; STW = sub-tropical water. Extended Abstracts

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limestones. Early Miocene sediments are also well exposed but not part of the current investigation. Research has been completed on the Cretaceous–Paleogene boundary sequence (Edwards 1966; Strong 1984; Brooks et al. 1986; Hollis and Strong 2003) and various fossil groups have been studied from this and surrounding sections over the past 150 years (Morgans et al. 2005). The Ashley Mudstone, a 60 m section of Early–Middle Eocene glauconitic mudstone, has been examined using TEX86, Mg/Ca, and δ18O to estimate sea temperatures from 50.7 to 46.5 Ma (Mangaorapan to Heretaungan Stages) (Hollis et al. 2009). Sea surface temperatures range from a maximum of 35°C in the lower section to a minimum of 25°C in the upper part of the section. Bottom water temperatures declined from 19–24°C in the lower part of the section to 16°C in the upper part of the section. MOERAKI-HAMPDEN The Moeraki-Hampden coastal section, 280 km to the south-west, has well exposed Paleocene–Early Eocene shelf to upper bathyal mudstones and siltstones, and mid to upper bathyal Early–Middle Eocene mudstones and muddy siltstones. Late Cretaceous strata crops out to the south (Shag Point) and Late Eocene– Oligocene strata crops out to the north (Kakanui River Mouth). A well-known and important molluscan locality (Beu and Maxwell 1990), and currently the stratotype for the Bortonian Stage (Cooper 2004), the sequence has also been examined for dinoflagellate cysts (Wilson 1985; Crouch and Brinkhuis 2005), calcareous nannoplankton (Edwards 1971a), and foraminifera (Finlay 1939a, b, c, d, 1940, 1947; Hornibrook 1961; Jenkins 1971). The Hampden Formation, a 120 m section of Middle Eocene silty glauconitic sandstone and sandy siltstone, contains very well-preserved foraminiferal faunas (Pearson and Burgess 2008) from which Mg/Ca and δ18O analyses, combined with TEX86 data from sediment, indicate sea surface temperatures of ca 23– 25°C (Burgess et al. 2008) during the Middle Eocene (42.7 Ma, Bortonian Stage). Figure 2 Straligrahic columns of Waipara and Hampden Beach sections.

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the

Mid-

Extended Abstracts


PALEOGEOGRAPHY The New Zealand micro-continent currently occupies a position from sub-tropical to subantarctic (30–55°S); during the Paleogene this position was considerably more southerly, ca 50–65°S in the earliest Eocene. Tectonic reconstructions for the Paleogene orientate the Canterbury Basin at ESE (Hollis et al. 2006, reconstruction at 55 Ma; Fig.1) A major geographic feature was the Chatham Rise to the east of the basin, possibly emergent in the Late Cretaceous and no deeper than shelf during the Paleogene. The Chatham Rise, and to a lesser extent the Campbell Plateau to the south, would have had a major effect on paleocurrents around the southern New Zealand sub-continent. Deposition of the midWaipara section may have been influenced more from the east, whereas the MoerakiHampden section appears to be more open to the south. Previous paleogeographic reconstructions (King et al. 1999; Hollis et al. 2006) place a large proportion of the western/northern Canterbury Basin at shelf depths (inner to mid) during the Paleocene–Eocene. Recent work on the mid-Waipara and Moeraki-Hampden sections indicates paleodepths of mid to upper bathyal from Early through Middle Eocene, suggesting a stronger influence from oceanic currents. STRATIGRAPHIC DRILLING A stratigraphic corehole, drilled at Lookout Bluff in 1968, near the top of the Hampden section penetrated 22.5 feet of gravels and 71.5 feet of late Eocene mudstone, tuff and sandstone (Edwards 1971b). Drilling was not overly successful and was abandoned before penetrating the targeted Bortonian (Middle Eocene) sediments. The core material was not accurately logged or curated and is of marginal use. Despite this, the potential for stratigraphic drilling at Hampden is very good with excellent access, a low dip (ca 4°), simple structure and a moderate thickness of 250–300 m of Late Paleocene to upper Eocene strata. Stratigraphic drilling at mid-Waipara may be more problematic. The section has a slightly steeper dip (ca 20°), is structurally simple and not overly thick (ca 300 m), and while getting

Extended Abstracts

drilling equipment into the valley could be difficult there are several possible access routes. Intriguingly, high Eocene sea temperatures from mid-Waipara and Hampden, along with well preserved fossil material, make these sections prime candidates for stratigraphic drilling. REFERENCES Brooks, R.R.; Strong, C.P.; Lee, J.; Orth, C.J.; Gilmore, J.S.; Ryan, D.E.; Holzbecher, J. 1986: Stratigraphic occurrences of iridium anomalies at four Cretaceous/Tertiary boundary sites in New Zealand. Geology 14: 727-729. Burgess, C.E.; Pearson, P.N.; Lear, C.H.; Morgans, H.E.G.; Handley, L.; Pancost, R.D.; Schouten, S. 2008: Middle Eocene climate cyclicity in the southern Pacific: Implications for global ice volume. Geology 36: 651-654. Cooper, R.A. ed. 2004: The New Zealand Geological Timescale. Institute of Geological and Nuclear Sciences Monograph 22: 284 p. Crouch, E.M.; Brinkuis, H. 2005: Environmental change across the Paleocene-Eocene transition from eastern New Zealand: A marine palynology approach. Marine Micropaleontology 56: 138-160. Edwards, A.R. 1966: Calcareous nanoplankton from the uppermost Cretaceous and lowermost Tertiary of the mid-Waipara section, South Island, New Zealand. New Zealand Journal of Geology and Geophysics 9: 481-490. Edwards, A.R. 1971a: A calcareous nannoplankton zonation of the New Zealand Paleogene. In Proceedings, 2nd Planktonic Conference, 1970, Rome, Italy: 381-419. Edwards, A.R. 1971b: Report on five stratigraphic holes drilled in North Otago, 1968. New Zealand Geological Survey Report 49. Finlay, H.J. 1939a: New Zealand Foraminifera: Key species in stratigraphy – No. 1. Transactions of the Royal Society of New Zealand 68: 504-533. Finlay, H.J. 1939b: New Zealand Foraminifera: The occurrence of Rzehakina, Hantkenina, Rotaliatina and Zeauvigerina. Transactions of the Royal Society of New Zealand 68: 534-543. Finlay, H.J. 1939c: New Zealand Foraminifera: Key species in stratigraphy – No. 2. Transactions of the Royal Society of New Zealand 69: 89-128.

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Finlay, H.J. 1939d: New Zealand Foraminifera: Key species in stratigraphy – No. 3. Transactions of the Royal Society of New Zealand 69: 309-329. Finlay, H.J. 1940: New Zealand Foraminifera: Key species in stratigraphy – No. 4. Transactions of the Royal Society of New Zealand 69: 448-472. Finlay, H.J. 1947: New Zealand Foraminifera: Key species in stratigraphy – No. 5. New Zealand Journal of Science and Technology 28: 259292.

Systematics and Distribution. New Zealand Geological Survey Paleontological Bulletin 34. Jenkins, D.G. 1971: New Zealand Cenozoic Planktonic Foraminifera. New Zealand Geological Survey Paleontological Bulletin 42. King, P.R.; Naish, T.R.; Browne, G.H.; Field, B.D.; Edbrooke, S.W. compilers 1999: Cretaceous to Recent sedimentary patterns in New Zealand. Institute of Geological and Nuclear Sciences Folio Series 1. Version 1999.1.

Hollis, C.J.; Strong, C.P. 2003: Biostratigraphic review of the Cretaceous/Tertiary boundary transition, mid-Waipara River section, North Canterbury, New Zealand. New Zealand Journal of Geology and Geophysics 46: 243253.

Morgans, H.E.G. in press: Late Paleocene–Middle Eocene stratigraphy, foraminiferal biostratigraphy and status of the Bortonian Stage lectostratotype at Moeraki-Hampden coastal section, eastern South Island, New Zealand. New Zealand Journal of Geology and Geophysics.

Hollis, C.J.; Crouch, E.M.; Dickens, G.R., 2006: How were Southwest Pacific pelagic ecosystems affected by extreme global warming during the initial Eocene thermal maximum? InterRad 11 and Triassic Stratigraphy Symposium, 2006, Wellington, NZ. Pp. 62.

Pearson, P.N.; Burgess, C.E. 2008: Foraminifer test preservation and diagenesis: comparison of high latitude Eocene sites. In Austin W.; James R. ed. Biogeochemical Controls on Palaeoceanographic Environmental Proxies. Geological Society of London, Special Publication 303. Pp. 59-72.

Hollis, C.J.; Handley, L.; Crouch, E.M.; Morgans, H.E.G.; Baker, J.A.; Creech, J.; Collins, K.S.; Gibbs, S.J.; Huber, M.; Schouten, S.; Zachos, J.C.; Pancost, R.D. 2009: Tropical sea temperatures in the high-latitude South Pacific during the Eocene. Geology 37: 99-102.

Strong, C.P. 1984: Cretaceous-Tertiary boundary, mid-Waipara River section, North Canterbury, New Zealand. New Zealand Journal of Geology and Geophysics 27: 231-234.

Hornibrook, N. de B. 1961: Tertiary foraminifera from Oamaru District (N.Z.): Part 1-

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Wilson, G.J. 1985: Dinoflagellate biostratigraphy of the Eocene Hampden section North Otago, New Zealand. New Zealand Geological Survey Record 8. Pp. 93-101.

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STABLE ISOTOPE RECORD OF SOIL CARBONATES FROM THE EOCENE– OLIGOCENE TRANSITION, BADLANDS NATIONAL PARK, SOUTH DAKOTA, USA Michelle R. Mullin and Richard H. Fluegeman Department of Geological Sciences, Ball State University, Muncie, IN 47306-0475, USA: [email protected]

The White River Group of the northern High Plains of the United States has produced a wealth of faunal and floral data across the Eocene – Oligocene transition. Sections in and around Badlands National Park in South Dakota are of particular interest as these have provided important paleontological data in defining the Chadronian and Orellan North American Land Mammal Ages (NALMAs) (Swisher and Prothero 1990; Ivany et al. 2003). The Chadronian – Orellan boundary is considered to represent the Eocene–Oligocene boundary in the western interior of North America (Swisher and Prothero 1990; Retallack 1992). Faunal and floral changes associated with Chadronian–Orellan interval have long been attributed to global climatic events associated with the GreenhouseIcehouse transition of the Eocene–Oligocene

Figure 1 Extended Abstracts

(Retallack 1992). The sections in the Badlands provide an opportunity to compare a response to global change in a temperate continental system with a response in a subtropical marine system (Gulf Coastal Plain, southeastern USA) and the world at large. Stable isotopes were obtained from paleosol carbonates across the Eocene–Oligocene boundary in four sections from three locations in Badlands National Park, Custer County, South Dakota (Fig. 1). Previous researchers have conducted paleomagnetic studies at these locations (Prothero and Swisher 1990; Prothero and Whittlesey 1998). Comparing the stratigraphy from the paleomagnetic studies to the stratigraphy of the samples collected for this research, the stable isotope samples can be constrained to Chron C13N-C13R transition,

Study area in South Dakota

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Oligocene (Fluegeman 2003; Miller et al. 2008). One of the interesting differences is in the actual numbers. The global curve shows an increase from approximately 2–3 ‰ (Zachos et al. 2001), while the regional Gulf Coastal Plain curve shows an increase from approximately -1 to 0 ‰ (Miller et al. 2008). The δ18O from the Badlands had an increase anywhere from -14 to -6 ‰. This allows for the interpretation of the δ18O curve from the mid-continent in terms of both temperature and aridity. Figure 2 Stable Isotopes at the Conata Picnic Area. Green line indicates Chadronian–Orellan boundary.

Figure 3 Stable Isotopes at the Conata Picnic Area. Green line indicates Chadronian–Orellan boundary.

which correlates to the Chadronian–Orellan NALMA. Results indicate climate stability through the Chadronian with wide fluctuations of oxygen isotope records in the uppermost Chadronian and lower Orellan (Figs 2 to 5). Records of carbon isotopes remain relatively stable through the studied interval. The results are consistent with changing climate associated with the Oi-1 glacial event in the early Oligocene but differ from Eocene– Oligocene patterns in the Gulf Coastal Plain. The Chadronian–Orellan δ18O curves seem to match the δ18O global curve from Zachos et al. (2001), which shows a slight decrease in δ18O in the upper Chadronian, followed by a marked increase at the boundary. This pattern is noted in three of the four measured sections sampled in the National Park. Oxygen isotope records across the Eocene–Oligocene boundary in the shallow marine record of the St. Stephens core of Alabama indicate a gradual cooling indicated by a gradual increase in δ18O through the late Eocene with a sharp decrease in temperatures associated with the early 96

While the signal of global climate change may be present in the Eocene–Oligocene strata δ18O of the Badlands, the δ13C appears to be very different. The Zachos et al. (2001) deep sea stable isotope curve also includes δ13C, which is anything but stable across the Eocene– Oligocene boundary. The global δ13C curve shows a slight decrease in the upper Eocene, followed by a marked increase at the Eocene– Oligocene boundary (Zachos et al. 2001). The δ13C found in the Badlands strata however does not follow this trend, essentially remaining stable, with a very slight increase across the Chadronian–Orellan boundary noticeable in only two of the measured sections. The stability in the Badlands δ13C demonstrates that local climate conditions may be more important in controlling the record of stable isotopes in paleosol carbonates in the region than global climate patterns. The stable isotope record from the paleosol carbonates in the Badlands seems to corroborate previous qualitative estimates of temperature and aridity changes in the region (Retallack 1983, 1992; Zanazzi et al. 2007). Paleosol type and degree of formation, as well as fossil assemblages and gaps have led previous researchers to determine that changes in water supply and mean annual temperature decreases occurred in the mid-continent, driving faunal migrations and extinctions (Retallack 1983, 1992; Zanazzi et al. 2007; Ivany 2003; Evans and Welzenbach 2000). It is now possible to show quantitatively that changes in temperature and aridity were indeed occurring, but not solely as a result of the global changes. Rather, local tectonic forces including the uplift of the Black Hills and regional volcanism played a role in creating a microclimate for the region. It is possible that the uplift west of the Badlands helped to create

Extended Abstracts


a rain-shadow, or that the relatively high ash content of the atmosphere in the mid-continent hampered moisture bearing cloud development. ACKNOWLEDGEMENTS The stable isotope research for the Badlands National Park, South Dakota, is ongoing as part of a Master’s Thesis. REFERENCES Evans, J.E.; Welzenbach, L.C. 2000: Lacustrine Limestones and Tufas in the Chadron Formation (Late Eocene), Badlands of South Dakota, U.S.A. Lake Basins through space and time: AAPG studies in Geology 46: 349-358. Fluegeman, R.H. 2003: Late Eocene-Early Oligocene benthic foraminifera in the Gulf Coastal Plain: regional vs. global influences. In Prothero, D.R.; Ivany, L.C.; Nesbitt, E.A. ed. From Greenhouse to Icehouse, the Marine Eocene-Oligocene Transition. New York, Columbia University Press. Pp. 283-293.

Figure 4 Stable Isotopes at Chamberlain Pass. Green line indicates Chadronian–Orellan boundary.

Ivany, L.C.; Nesbitt, E.A.; Prothero, D. 2003: The Marine Eocene-Oligocene Transition, a synthesis. In Prothero, D.R.; Ivany, L.C.; Nesbitt, E.A. ed. From Greenhouse to Icehouse, the Marine Eocene-Oligocene Transition. New York, Columbia University Press. Pp. 522-534. Miller, K.G.; Browning, J.V.; Aubry, M.-P.; Wade, B.S. 2008: Eocene-Oligocene global climate and sea-level changes: St. Stephens Quarry, Alabama. Geological Society of America Bulletin 120: 34-53. Prothero, D.R.; Whittlesey, K.E. 1998: Magnetic stratigraphy and biostratigraphy of the Orellan and Whitneyan land-mammal “ages” in the White River Group. In Terry, D.O.; LaGarry, H.E.; Hunt, R.M. Jr. ed. Geological Society of America Special Paper 325 - Depositional Environments, Lithostratigraphy, and Biostratigraphy of the White River and Arikaree Groups (Late Eocene to Early Miocene, North America). Boulder, Geological Society of America, Inc. Pp. 39-62. Retallack, G.J. 1983: A Paleopedological approach to the interpretation of terrestrial sedimentary rocks: The mid-Tertiary fossil soils of Badlands National Park, South Dakota. Geological Society of America Bulletin 94: 823-840.

Figure 5 Stable Isotopes at Dillon Pass. Green line indicates Chadronian–Orellan boundary. Extended Abstracts

Retallack, G.J. 1992: Paleosols and changes in climate and vegetation across the

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Eocene/Oligocene Boundary. In Prothero, D.R.; Berggren, W.A. ed. Eocene-Oligocene Climatic and Biotic Evolution. Princeton, New Jersey, Princeton University Press. Pp. 382398. Swisher, C.C. III; Prothero, D.R. 1990: Single Crystal 40Ar/39Ar dating of the EoceneOligocene Transition in North America. Science 249: 760-762.

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Zachos, J.C.; Pagani, M.; Sloan, L.; Thomas, E.; Billups, K. 2001: Trends, Rhythms, and Aberrations in Global Climate 65 Ma to Present. Science 292: 686-693 Zanazzi, A.; Kohn, M.J.; MacFadden, B.J.; Dennis O.T. Jr. 2007: Large Temperature drop across the Eocene-Oligocene transition in central North America. Nature 445: 639-642.

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EOCENE OCEAN TEMPERATURES AND CLIMATE GRADIENTS Paul N. Pearson School of Earth and Ocean Sciences, Cardiff University, Cardiff CF10 3YE, UK: [email protected]

The oceans and atmosphere distribute heat from the tropics towards the poles, defining broad climatic belts. The modern world has marked gradients in open ocean sea surface temperatures, which range from up to 30°C in the tropics to freezing near the poles, and are quite consistent in all three major ocean basins and in both hemispheres. A first order challenge that we have in understanding times in the geological past is to reconstruct the sea surface temperature gradient, which may have been very different from today with profound consequences for global and regional climate. This gradient will be influenced by both the global average temperature (affected in turn by the solar constant, the albedo and the greenhouse effect), local effects such as the presence of continents, orogens, ice sheets and albedo variations, and the efficiency of heat transport systems on the planet (related to ocean and atmospheric circulation patterns). Data from the rock record is sparse and limited, but can be compared against climate model outputs to help investigate the likely processes of heat transport and infer the climatic boundary conditions (e.g. greenhouse effect) required to explain the data. At the same time, past climate states provide challenging tests for the climate models themselves. Widespread fossil and other geological evidence has suggested that the Eocene was for the most part much warmer than present with little or no ice for most of the time. Historically the first and still the most important geochemical proxy for Eocene seawater temperatures is oxygen isotope analysis of plankton shells. Initial data from deep-sea drilling were challenging, however, especially in the tropics, where temperatures from planktonic foraminifera were commonly reconstructed as significantly cooler than modern (various studies following the paradigm of Shackleton and Boersma 1981). This proved difficult to simulate with climate models and conflicted with inferences based on the known temperature tolerances of a range of biota (Adams et al. 1990). It is now clear that

Extended Abstracts

the vast majority of foraminifer shells from carbonate oozes and chalks of this age have been pervasively recrystallized during shallow burial, which likely explains the anomalous cool temperatures. Renewed focus on unrecrystallized foraminifera from hemipelagic clay-rich facies in several key localities, supported by trace element and organic geochemical data, now indicates much warmer temperatures over a range of latitudes than previously estimated. The same diagenetic problem affects oxygen isotope data from benthic foraminifera. However compilations of benthic foraminiferal oxygen isotope data from deep-sea cores (e.g. Zachos et al. 2008) remain crucial for understanding the long-term trends in Cenozoic climate. Even though deep-sea benthic foraminifera also tend to be recrystallized, the temperature bias is likely less pronounced than the planktonic records because recrystallization happens in similar cool sea-floor conditions to the original habitat of the foraminifera. The deep-sea benthic compilation reveals the episodic cooling and ice growth that happened over the Cenozoic, including a pronounced cooling trend from the Early Eocene to the Oligocene which culminated in a major shift in climate over the Eocene–Oligocene transition (Zachos et al. 2008). The benthic compilation is not, however, a ‘global climate curve’ as is sometimes represented, insofar as it reflects only the cold end of the climate gradient. This is because oceanic bottom waters are thought to be mostly derived from the cold high latitudes, and continental ice growth also tends to occur at high latitudes. A major goal is therefore to determine oxygen isotope paleotemperatures from tropical planktonic foraminifers to detect climate trends at the warm end of the global spectrum. This will reveal how climate gradients have changed through time and allow a better understanding of the likely forcing of global climate trends. It will also permit quantitative reconstructions of the average global temperature which cannot be derived from high latitude records alone.

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Well-preserved Eocene planktonic foraminifera are now known from a growing list of locations (e.g. New Jersey, US Gulf Coast, Mexico, New Zealand, the north Atlantic, the Adriatic, India, Java, Tanzania and Mozambique). The most complete and consistent tropical record for the Paleogene is from multiple drill sites and outcrop sections in Tanzania (e.g. Pearson et al. 2007). Moreover, unlike some of the other clay-rich sites, the Tanzanian sediments were deposited in relatively deep oceanic water (e.g. Bown et al. 2008) and so are probably less affected by local environmental variability. Although so far at a relatively low temporal resolution, the Tanzanian record shows relatively warm and stable background climate through the Eocene, with reconstructed sea-surface temperatures typically in the low 30s (which is several degrees warmer than modern). This is apart from the so-called ‘hyperthermals’ and other known transient warm phases when temperatures were presumably considerably warmer than this. Also, comparisons with upper Eocene Paleotemperatures with drillcores in Java (Pearson and Coxall unpublished data) suggest that, as expected, the Indian Ocean off Tanzania was not the warmest part of the world, it being a little warmer in Java.

The Tanzanian oxygen isotope temperatures are supported by trace element and organic paleotempertaure estimates. There is a good fit between the isotopes and the TEX86 proxy (Pearson et al. 2007) but only if a relatively cool calibration for TEX86 is used. More recent ‘warm’ calibrations (e.g. Kim et al. 2008) would indicate sea surface temperatures approaching 40°C in Tanzania, which does not fit the isotope data, and such warm conditions are also difficult to square with the known distributions of fossil biota. It is likely therefore that the ‘warm’ calibrations for TEX86 overestimate temperatures in the Paleogene. This is feasible because there is no reason why a calibration based on modern crenarcheota, which do not today experience such consistently warm open ocean conditions, should necessarily apply quantitatively in the past. An evolutionary response by the crenarcheota to changing global conditions cannot be ruled out. Given this, it may be advisable to develop a ‘paleo-calibration’ for TEX86 based on locations with well-preserved planktonic foraminifera on which stable isotope and trace element geochemistry can be conducted. The beauty of the TEX86 proxy is that it can then be used in a much wider variety of geological settings than the stable isotopes.

It can be concluded from this that the global average temperature was likely much warmer in the Eocene than has been estimated from the benthic isotope data alone. It is likely that very large portions of the tropical and sub-tropical ocean were consistently warmer than 30°C for much of the year throughout the Eocene. This has interesting implications for the evolution of the tropical biota, for weather (e.g. tropical storm frequency and intensity) and climate.

REFERENCES

One feature of the Tanzanian data is that the long-term cooling trend in the Eocene occurred mainly at high latitudes, whereas the tropics remained very warm throughout the epoch (Pearson et al. 2007). In other words, the global climate gradient increased markedly from the early to late Eocene. This may be consistent with a gradual reduction in greenhouse gas forcing, the effects of which were amplified at the poles, culminating with a threshold at the Eocene–Oligocene transition caused by explosive ice sheet growth (e.g. Deconto and Pollard 2003)

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Adams, C.G.; Lee, D.E.; Rosen, B.R. 1990: Conflicting isotopic and biotic evidence for tropical sea-surface temperatures during the Tertiary Palaeogeography, Palaeoclimatology, Palaeoecology 77: 289-313. Bown, P.R.; Dunkley Jones, T.; Young, J.R.; Randell, R. 2009: A Paleogene record of extant lower photic zone calcareous nannoplankton. Palaeontology 52: 457-469. Deconto, R.M.; Pollard, D. 2003: Rapid Cenozoic glaciation of Antarctica induced by decliningatmospheric CO2. Nature 421: 245-249. Kim, J.-H.; Schouten, S.; Hopmans, E.C.; Donner, B.; Sinninghe Damste, J.S. 2008: Globalsediment core-top calibration of the Tex86 aleothermometer in the ocean. Geochimica et Cosmochimica Acta 72: 1154-1173. Pearson, P.N.; van Dongen, B.E.; Nicholas, C.J.; Pancost, R.D.; Schouten, S.; Singano, J.M; Wade, B.S. 2007: Stable warm tropical climate through the Eocene epoch. Geology 35: 211214.

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Shackleton, N.J.; Boersma, A. 1981: The climate of the Eocene ocean. Journal of the Geological Society of London 138: 153-157.

Extended Abstracts

Zachos, J.C.; Dickens, G.R.; Zeebe, R.E. 2008: An early Cenozoic perspective ongreenhouse warming and carbon cycle dynamics. Nature 451: 279-283.

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RECONSTRUCTING THE LATE PALEOCENE–EARLY EOCENE CONTINENTAL PALEOSURFACE IN AND AROUND THE PARIS AND ADJACENT BASINS: NEW INSIGHTS FOR PALEOGEOGRAPHIC, GEODYNAMIC AND CLIMATIC STUDIES F. Quesnel1, C. Dupuis2, J. Yans3, C. Ricordel-Prognon1, S. Rad1, J.-Y. Storme3, F. Barbier3, E. Roche4, C. Bourdillon5, T. Smith6 and P. Iacumin7 1

BRGM (French Geological Survey), GEO/G2R, 1, Avenue Claude Guillemin, BP 36009, 45060 – Orléans Cedex 2, France: [email protected]; 2Polytechnics, GFA, rue de Houdain 9, 7000 – Mons, Belgium; 3UCLNamur, FUNDP., rue de Bruxelles, 61, 5000 – Namur, Belgium; 4ULg, Paléontologie végétale, Sart Tilman, B18/P40, 7000 – Liège, Belgium; 5ERADATA, 5, Allée des Magnolias, 72100 Le Mans, France; 6IRSNB, Département de Paléontologie, 29 rue Vautier, B-1000 Bruxelles, Belgium; 7Universita' di Parma. Dipartimento di Scienze della Terra Via Usberti 157/A. 43100 – Parma, Italy.

INTRODUCTION, AIMS AND METHODS The geological archive records "hyperthermal" crises, along with their consequences for the biotic and physical environment. Among these, the Paleocene–Eocene Thermal Maximum (PETM) is considered the closest analogue to the current climate crisis, due to its global character and speed at which the CO2 rate and average temperatures increased (Higgins and Schrag 2006; Zachos et al. 2001, 2008). Some 55.8 Ma, it affected the Earth for a period of almost 200 k.y. (Röhl et al. 2000; Westerhold et al. 2007), and continental and marine paleoenvironments were marked by a negative δ13C anomaly coinciding with a negative δ18O anomaly indicative of a notable temperature rise (3–8°C). We are developing a multi-disciplinary study in the coastal to continental paleoenvironment of the Sparnacian facies preserved within the Paris and adjacent basins, and also on the neighbouring basements that are sites of contemporary weathering. We aim to integrate the upstream to downstream sequence of diversified paleoenvironments, their landscapes and ecosystems, in order to assess the impact of the PETM climate crisis on each and the whole sequence. The preliminary work described here has involved: •

compiling, reviewing and validating all available evidence;

•

making new field observations in poorly investigated areas;

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•

studying occurrences of paleo-weathering and continental deposits, dating these using a variety of tools (e.g. litho, bio and chemo-stratigraphy, paleomagnetism, Ar/Ar geochronology), and relating them to the P–E continental paleosurface;

•

organizing all the data in a GIS Database, using these to digitally reconstruct the current geometry of this paleosurface at 1:1 000 000 scale, then drawing the geological cross sections through the studied basins and their borders;

•

reconstructing the continental paleogeography of the Paris and adjacent basins and their surroundings during this interval (Fig. 1). FIRST RESULTS

Historically, the Paris and adjacent basins are the cradle of stratigraphy, where the notion of "Sparnacian" took shape (Dollfus 1880; see the detailed lithostratigraphy in Aubry et al. 2005). The Sparnacian facies, mainly continental to coastal, often display paleo-weathering features. Near the Mesozoic cover and older basement surrounding those basins, many fluviatile sands and conglomerate units seal or incise older thick kaolinitic weathering profiles. Sedimentological and stratigraphic studies and the mapping of those deposits show them associated with a major unconformity, spanning the uplifted areas to the lowlands of the shallow basins where they incise marine formations of Late Thanetian age. Almost all these paleo-weathering profiles and fluvial deposits are oxidized and/or leached and silicified in and around the studied

Extended Abstracts


Figure 1 Paleogeographic map of NW Europe and North Atlantic around the Paleocene–Eocene boundary, modified after Ziegler (1988), and showing the continental facies and paleo-weathering types on emerged lands

basins, with pedogenic silcretes upstream andquartzitic silcretes on the leached fluviatile deposits of the lowlands, originally rich in lignite and pyrite. These silcretes are probably the most striking geological markers of the P– E paleo-weathering, and were long known to Extended Abstracts

geologists and geomorphologists who tried to map the “Eocene continental paleosurface”. They are often ascribed to Late “Landenian” to the north of the Paris-Belgian Basins and correlated to Sparnacian continental deposits to the south, west and east of the Paris Basin 103


(Dupuis 1979; Thiry 1981; Quesnel 1997). The silcretes also crop out in England — the pedogenic facies close to the Devon, and the quartzitic facies, termed Sarsenstone, which is more widespread around the London and Hampshire Basins, and locally found in situ in the Reading Beds. Some silcretes are more oxidized, giving ochreous and red sandstones (e.g. “Pays d’Ouche” in Normandy and Grandglise Sandstones in Belgium). Above the upstream paleo-weathering profiles, the silcretes and oxidized sandstones are almost never overlain by other formations, preventing any precise dating. Yet they are sealed by lacustrine limestones of Lutetian to Bartonian age in a few small grabens in Normandy, Perche, and on the Beauce margins. In the best preserved successions downstream, these weathering profiles lie above, or are developed upon, marine or continental upper Thanetian formations and are overlain by lower Ypresian marine formations, in England, northern France-Belgium and Upper Normandy. Locally they are the stratigraphic equivalents to fluvial sands containing lignitic units where the PETM has been recorded (e.g. Belgium (Steurbaut et al. 2003) and Northern France (Magioncalda unpublished data; Quesnel 2006; Storme et al. unpublished data)). The paleomagnetic ages obtained from the silcretes and oxidized sandstones from the “Pays d’Ouche” and Grandglise Sandstones indicate a paleoweathering episode around the P–E boundary (Ricordel-Prognon et al. unpublished data). The 39Ar-40Ar dating of the supergene Mn oxides from the Morialmé weathering profile formed on the Ardenne basement also gives an age around the P–E boundary (Barbier et al. unpublished data). Finally, the quartzitic and oxidized silcretes (Sarsenstones and Landenian Sandstones) are reworked in a few small very well rounded pebbles that accompany flint pebbles within the coastal Blackheath and Oldhaven Formations (Early Ypresian) in England (Stamp 1921) and the “Conglomérat à galets avellanaires”, their stratigraphic equivalent in Upper Normandy and Avesnois (Quesnel 2006 unpublished data). All this data implies that the paleo-weathering that produced the silcretes and oxidized sandstones occured after the Late Thanetian and prior to the lower Ypresian transgression. Additionally, the kaolinitic weathering profiles were clearly

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formed earlier, and probably under wetter conditions, than the silcretes, mainly during the Early Cretaceous on the old basement and Jurassic formations (Yans et al. 2003; Thiry et al. 2006; Ricordel, 2007; Théveniaut et al. 2007) and during the Paleocene on the Chalk (Quesnel et al. 2007). DISCUSSION The widespread occurrence of a peculiar paleoweathering in and around the Paris and adjacent basins is present around the P–E boundary. We have recognised pedogenic, quartzitic and oxidized silcretes in Limbourg, Thiérache, Upper Lorraine, Luxembourg, Brittany, Touraine, northern Aquitaine Basin, around the Morvan and the Massif Central, on the Bresse edge and they are also described in Germany in the Saale-Elbe Basin (Eissmann 2002), Eifel (Löhnertz 2003 pers. comm.) and in Hesse (Thiry 2003 pers. comm.). The processes involved in generating these types of silcretes are relatively well known in the Paris and adjacent basins and their borders (Dupuis and Steurbaut 1987; Thiry 1999) and some appear to be closely linked to effects in relation with a climate crisis such as the PETM (i.e. marked alternations of flooding, inducing clay deposition in pores and soil cracks, followed by dry phases saturating the ground solutions; acid drainage of highly organic and pyrite-rich sediments, destabilization of kaolinite). The silica may have been provided by the weathering of the quartz sands and of the flints of the Clay-with-flints (Cwf) profiles formed during the Paleocene at the expense of the Chalk, which largely covered the area at the end of the Cretaceous (Quesnel et al. 2007). Other types of paleo-weathering, such as variegated clay, oxide nodules, gley and pseudogley features formed in the lacustrine and fluviatile formations along the P–E continental paleosurface on the borders of the Sparnacian basin (Buurman 1980; Thiry 1981; Laurain and Meyer 1986). These probably formed in a less well drained paleoenvironment in the clayey lowlands. Sparse calcretes also occur in similar P–E formations south and east of the Paris Basin. These are much more common in the Languedoc, Provence and Pyrenees Garumnian facies (i.e. the Upper Paleocene and Lower Eocene continental formations; Plaziat et al.

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1987; Cojan et al. 2000; Schmitz and Pujalte 2003, 2007). Here, they also probably formed in alternating, strongly seasonal, wet and dry conditions, but in a less acidic paleoenvironment due to carbonate input from the surrounding landscape. We use the silcretes and all other contemporaneous geological markers to reconstruct the P–E continental paleosurface and to draw its features on the paleogeographic map presented in Figure 1. Added to appropriate climatic conditions, the shaping of this continental paleosurface was probably triggered at the end of the Paleocene by lithospheric deformation related to the first steps of the Pyreneo-alpine orogeny and to the rifting and volcanic activity of the NAIP, and also accompanied by significant sea level variations. Once the continental paleogeography of the P– E transition in the Paris and adjacent basins has been established, we aim to: •

•

•

refine the litho-, bio- and chemostratigraphic record of the sedimentary units from coastal to continental paleoenvironments; study and date other weathering profiles along the paleotoposequence from the uplands to the lowlands; and study the evolution of several proxies to determine the impact of the PETM climatic crisis on the processes (and their intensities) affecting the various compartments of the P–E continental paleosurface.

Furthermore, this regional study may be directly used in improving the tools for simulating the landscape evolution and the Earth's climate in paleoenvironmental modeling. REFERENCES Aubry, M.-P.; Thiry M.; Dupuis, C; Berggren, W.A. 2005: The Sparnacian deposits of the Paris Basin: Part I. A litho stratigraphic classification. Stratigraphy 2/1: 65-100. Buurman, P. 1980: Palaeosols in the reading Beds (Paleocene) of Alum Bay, Isle of Wight, UK. Sedimentology 27: 593-606.

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Cojan, I.; Moreau, M.-G; Stott, L.D. 2000: Stable Carbon isotope stratigraphy of the Paleogene pedogenic series of southern France as a basis for continental-marine correlation. Geology 28: 259-262. Dollfus, G. 1880: Essai sur l’extension des terrains tertiaires dans le Bassin anglo-parisien. Bulletin Societé géologique Normandie VI: 584–605. Dupuis, C. 1979: Esquisse paléogéographique du Nord et du Nord-Ouest du Bassin de Paris au Paléocène et à l'Eocène inférieur. Incidences structurales. C. R. Acad. Sci. Paris 288: 15871590. Dupuis, C.; Steurbaut, E. marins (NP8, NP9) et et stromatolites dans entre Criel et le Normandie). Annales Nord CV: 233-242.

1987: Altérites, sables fluviatiles, silicification le Paléocène supérieur Cap d'Ailly (HauteSociété Géologique du

Eissmann, L. 2002: Tertiary geology of the SaaleElbe region. Quaternary Science Reviews 21: 1245-1275. Higgins, J.A.; Schrag, D.P. 2006: Beyond methane: Towards a theory for the Paleocene-Eocene Thermal Maximum. Earth and Planetary Science Letters 245: 523-537. Laurain, M.; Meyer, R. 1986: Stratigraphie et Paléogéographie de la base du Paléogène champenois. Géologie de la France 1: 103123. Plaziat, J.C.; Freytet, P.; Marec, P. 1987: Sédimentation molassique et paléopédogenèse en Languedoc, les dépôts fluviatiles, lacustres et palustres du Maastrichtien au Bartonien. Field Trip Guide, 16-18 September 1987. ASF Publications. 127 p. Quesnel, F. 1997: Cartographie numérique en géologie de surface - Application aux altérites à silex de l'ouest du Bassin de Paris. PhD thesis, University of Rouen. Document du BRGM 263. 268 p and 160 p of annexes. Quesnel, F. 2006: Méso-Cénozoïque, Crétacé et Cénozoïque, 3 geological maps and 87 p of explanatory notes, In Lacquement, F.; Quesnel, F.; Mansy, J.L.; Moulouel, H.; Vinchon, C.; Gateau, C. ed. La Géologie du territoire de l’Avesnois, Système d’information géologique, DVD, BRGM/RP-55465-FR. Quesnel, F.; Bourdillon, C.; and coll. 2007: Résidus à silex de l’Ouest du bassin de Paris (Normandie et Perche), Du terrain à la typologie des faciès, Biostratigraphie des silex résiduels, Cartographie numérique, Modélisation géométrique, évolution des profils d’altération sur le Crétacé supérieur,

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Field Trip Guide, 26 October 2007. ASF Publications 61. 88 p. Ricordel, C. 2007: Datations par paléomagnétisme des paléoaltérations du Massif Central et de ces bordures : implications géodynamiques. PhD thesis memoir, Ecole des Mines de Paris. 172 p. Röhl, U.; Bralower, T.; J., Norris, R. D.; Wefer, G. 2000: New chronology for the late Paleocene thermal maximum and its environmental implications. Geology 28: 927-930. Schmitz, B.; Pujalte, V. 2003: Sea-level, humidity, and land-erosion records across the initial Eocene thermal maximum from a continentalmarine transect in northern Spain. Geology 31: 689-692. Schmitz, B.; Pujalte, V. 2007: Abrupt increase in seasonal extreme precipitation at the Paleocene-Eocene boundary. Geology 35: 215218. Stamp, L.D. 1921: On the beds at the base of the Ypresian (London Clay) in the Anglo-FrancoBelgian basin. Proc. Geol. Assoc. XXXII: 57110. Steurbaut, E.; Magioncalda, R.; Dupuis, C.; Van Simaeys, S.; Roche, M.; Roche, E. 2003: Palynology, paleoenvironments and organic carbon isotope evolution in lagoonal Paleocene-Eocene boundary settings in North Belgium. In Wing, S.L.; Gingerich, P.D.; Schmitz,B.; Thomas E. ed. Causes and consequences of globally warm climates in the early Paleogene, Geological Society of America Special Paper 369. Pp. 291-317. Théveniaut, H.; Quesnel, F.; Wyns, R.; Hugues, G. 2007: Paleomagnetic dating of the “Borne de Fer” ferricrete (NE France): Lower Cretaceous continental weathering. Palaeogeography, Palaeoclimatology, Palaeoecology 253: 271279. Thiry, M. 1981: Sédimentation continentale et altérations associées: calcitisations, ferruginisations et silicifications. Les Argiles Plastiques du Bassin de Paris, Sciences Géologiques Memoirs 64. 173 p.

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Thiry, M. 1999: Diversity of continental silicification features: examples from the Cenozoic deposits in the Paris Basin and neighbouring basement. In Thiry, M.; SimonCoinçon, R. ed, Palaeoweathering, Palaeosurfaces and Related continental deposits. IAS Special Publication 27, Blackwell Science. Pp. 87-127. Thiry, M.; Quesnel, F.; Yans, J.; Wyns, R.; Vergari, A.; Théveniaut, H.; Simon-Coinçon, R.; Moreau, M.-G.; Giot, D.; Dupuis, C.; Bruxelles, L.; Barbarand, J.; Baele, J.-M. 2006: Continental France and Belgium during the Early Cretaceous: Palaeoweathering and palaeolandscapes. Bulletin de la Societe Géologique de France 177: 155-175. Westerhold, T.; Röhl, U.; Laskar, J.; Bowles, J.; Raffi, I.; Lourens, L.J.; Zachos, J.C. 2007: On the duration of magnetochrons C24r and C25n and the timing of early Eocene global warming events: implications from the Ocean Drilling Program Leg 208 Walvis Ridge depth transect. Paleoceanography 22: PA2201. doi:10.1029/2006PA001322. Yans, Y.; Quesnel, F.; Dupuis, C. 2003: Mesocenozoic paleoweathering of the Haute-Lesse area (Ardenne-Belgium). Field trip guides of the Special Conference on Paleoweathering and Paleosurfaces in the Ardennes-Eiffel Region, Bettborn/Preizerdaul, Luxembourg, Field Trip 1, may 16, 2003, morning. Géologie de la France 2003-1: 3-10. Zachos, J.C.; Pagani, M.; Sloan, L., Thomas, E., Billups, K. 2001: Trends, rhythms and aberrations in global climate 65 Ma to Present. Science 292: 686-693. Zachos, J.C.; Dickens, J.R.; Zeebe R.E., 2008: An early Cenozoic perspective on greenhouse warming and carbon-cycle dynamics. Nature 451: 279-283. Ziegler, P.A. 1988: Evolution of the Arctic-North Atlantic and the Western Tethys. American Association of Petroleum Geologists Memoir 43. 198 p, 30 pl.

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PALEOCENE–EOCENE AGE AND PALEOENVIRONMENTS, NORTH-EASTERN MARGIN OF AUSTRALIA – WORK IN PROGRESS Patrick G. Quilty School of Earth Sciences, University of Tasmania, Private Bag 79, Hobart, Tasmania 7007 Australia: [email protected]

INTRODUCTION Pujalte et al. (2009) showed that detailed study of the Paleocene–Eocene Thermal Maximum (PETM) was heavily biased towards the Mediterranean region and that there has been no study over a vast area. This paper summarises knowledge of Paleocene–Eocene sections following recent dredging along the coast off north-eastern Australia. It provides hope that continuous Paleocene–Eocene sections exist to add information from a region evolving from mid to low latitudes during this interval. Little is known of the Paleogene marine faunas and paleoenvironments along the eastern margin of Australia. Onshore marine sections are lacking and thus data are gathered from chance discoveries and dedicated marine

research cruises that have been conducted to ascertain the tectonic evolution of continental fragments and seamounts. Deep Sea Drilling Project (DSDP, Leg 21) and Ocean Drilling Program (ODP, Leg 133) have drilled in the region but no continuous Paleocene–Eocene transition sections were encountered. The Paleogene material reported here was recovered during a series of research cruises conducted by Geoscience Australia, particularly cruises 270 and 274 to Mellish Rise and Kenn Plateaus (Exon et al. 2006a,b) and 295 to Marion Plateau (Fig.1). The area is about 270 000 km2. The overall results have been summarised in Hoffmann et al. (2008) and that is used as the background to this paper. Ages identified are concentrated in the Late Paleocene (P4), Mid-Eocene (P10/11) and Late Eocene (P15/16). These ages are well recorded from western and southern Australia and thus the Australia-wide sedimentation cycle pattern seems to be reflected here (Fig.2). At ca 120 Ma, the regional tectonic regime changed from compressional to extensional as Lord Howe Rise and New Caledonia-Norfolk Ridge formed in association with opening of the Tasman Sea to the south. The margin began seriously to break up at about K–T (65 Ma) with formation of the Coral Sea, Cato and northern Tasman Sea Basins, and formation of thinned continental fragments. Spreading ceased at ca 52 Ma. MELLISH RISE Mellish Rise is the most heavily sampled feature and can be separated into Northern and Southern Mellish Rises separated by a narrow northeast-southwest oriented trough. Most samples contain identifiable foraminiferal faunas that allow age determination.

Figure 1 Location studied for this paper Extended Abstracts

off

Queensland

of

samples

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Figure 2

Summary of dating and lithology of studied samples.

Dredge DR 14 yielded two datable faunas of which Lithology G contains a beautiful fauna from the latter part of Middle Eocene, P14. Deposition appears to have been as deep-sea ooze but could be a continental slope deposit. Lithology F1.1 yielded an excellent Late Paleocene, P4a/b, fauna and deposition occurred under outer shelf or upper slope conditions. DR15 Lithology C contains a magnificent benthic and planktonic fauna of Late Paleocene, P4, probably 4a/b. The environment of deposition appears to be mid shelf. DR13 Lithology D1.1 is foram/nanno ooze with a Late Paleocene (P4c) fauna. DR24 Lithology A from Mellish Reef is dominantly of Late Eocene Indo-Pacific large foraminifera such as Asterocyclina/Discocyclina and Nummulites. The locality is now some 2000 m below where it was deposited and if in situ, the area has sunk some 2000 m over the last 40 m.y.. DR25 Lithology B includes Asterocyclina sp. (probably A. incisuricamerata), Crespinina sp. and Gypsina cf. disca of undifferentiated Eocene but likely to be Mid- to Late Eocene. Deposition occurred within the photic zone and

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the locality may have sunk some 2500 m since deposition. DR31 Lithology A is grey marl with abundant moderately well-preserved foraminifera consistent with a Late Eocene (P15) age. Deposition occurred in a mid-shelf environment off a coast where there was fine terrigenous sediment input, in turn suggesting exposed land nearby. Lithology D has an excellent mixed large/planktonic fauna of earliest Late Eocene, P15, age and the environment of deposition was mid to outer shelf depths. DR40 was a major dredge haul with diverse lithologies. All Lithology A rocks have abundant large foraminifera and a few planktonic species. They are highly fragmented attesting to a high-energy, shallow-water environment within the photic zone. The age is Middle Eocene, probably in the upper part, P14/15. Lithology B fauna includes Asterocyclina sp. and Nummulites sp. (to ca 20 mm diameter) and is thus Middle-Late Eocene. Deposition occurred in shallow water in photic zone. Lithology J is now a high-porosity,

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weakly-cemented biocalcarenite but identification is difficult. Planktonic species are quite common but low diversity and lacking diagnostic forms. The fauna includes Discocyclina/Asterocyclina sp., Lepidocyclina sp. and common shallow-water benthic species. The age is Mid-Late Eocene as with other samples in this dredge haul. The environment of deposition was a shallowwater, high-energy carbonate sand with no terrigenous input. DR43 Lithology A contains Biplanispira cf. mirabilis, Lepidocyclina sp., possibly Discocyclina and Fabiania sp. allowing correlation with the Late Eocene age (Tertiary b in the Indo-Pacific Letter classification) or P14–P17. This is the first record of Biplanispira in Australasia. Lithology C is of extremely well sorted biocalcarenite. Very little of the fauna is identifiable but one of the large specimens is either Biplanispira or Pellatispira. The age is likely to be Late Eocene, Tertiary b in the Indo-Pacific Letter classification (P14-P17). DR19 Lithology A4.2 contains shapes of foraminifera such as G. acuta/formosa/marginodentata and other species similar to G. aequa, Globanomalina wilcoxensis, Chiloguembelina sp. and Catapsydrax sp. A of Late Paleocene–Early Eocene age. DR44 Lithology A samples are highly altered and calcareous nannofossils (Findlay in Hoffmann et al. 2008) suggesting a Paleocene age, probably P4a/b. KENN PLATEAU 270/DR06 Lithology A1.1 consists of white, largely recrystallised ooze containing Truncorotaloides rohri, Globorotalia broedemanni, Globigerina officinalis, G. frontosa, G. yeguaensis, “Globigerapsis index” and Pseudogloboquadrina primitiva. The main body of residue belongs to the earlier part of Middle Eocene (P10/11) and probably was deposited in mid shelf depth. Lithology B1 yielded a small residue with pristine foraminifera, including Pseudohastigerina micra, Catapsydrax unicavus, C. africanus, Chiloguembelina cubensis, Streptochilus martini, Globorotalia cerroazulensis, G.

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cocoaensis, Subbotina gortanii, Globigerina ‘winkleri’, G. corpulenta, G. linaperta and G. angiporoides. The age is Late Eocene, P14-16, and deposition seems to have occurred at ooze depths. 274/DR07 Lithology B1.1 consists of white calcarenite with abundant large foraminifera and Globigerinatheka index tropicalis, G. index s.l., G. mexicana barri, Globigerina yeguaensis, Catapsydrax unicavus, C. dissimilis and large forms such as Asterocyclina incisuricamerata of Eocene, Late P14/Early P15, Tertiary a age. Deposition occurred within the photic zone in association with seagrass beds below wave base. The site has thus sunk at least 1500 m since deposition. MARION PLATEAU Cruise 295/DR03 Lithology B2 is of highly calcareous greensand with small gastropod fragments and few foraminifera both benthic and planktonic. The foraminiferal fauna is from near the Middle-Late Eocene boundary, P14-P15, and deposition occurred in mid-shelf depths. Abundant glauconite suggests that the Middle-Late Eocene age applies also to the rest of the highly glauconitic sediments from the same dredge haul. CONCLUSIONS Samples have yielded Paleocene and Eocene ages. Several include noteworthy species that have not been recorded from this area before, although many are known from elsewhere in the Indo-Pacific region and provide a means of reconstructing biogeography. They include those from a variety of depositional environments, many with large foraminifera such as Operculina, Discocyclina, Asterocyclina, Heterostegina, and others that, while sometimes classified as small, are studied using the same techniques as traditional ‘large’ foraminifera. The large species generally are small (mostly to some 4.5 mm but Nummulites to 20+ mm). Several samples contain excellent foraminiferal and ostracod faunas crying out for further study and documentation. Paleocene faunas are of small species with good to excellent planktonic faunas. The absence of possible Indo-Pacific forms may

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suggest that conditions were cooler than in the Eocene. If all sections suffered the same subsidence history, it seems that the environment was shallow-water, sub-tropical to tropical throughout the Eocene. These records indicate that the Paleocene– Eocene transition should be preserved in the region, possibly most likely in troughs between the continental blocks. This may form part of justification of deep-ocean drilling in the area. Marion Plateau samples contain Mid-Late Eocene (P14/15) planktonic faunas of tropical, mid-shelf aspect. They accumulated on a fragment of thinned continental crust, an extension of the Australian mainland. Several sectors produced no specific results and further sampling is needed to help elucidate the history of Kelso Basin, Northern Mellish Rise, Selfridge Rise, and the mystery of a low birefringence material from Mellish Rise. Radiolaria may help decipher some of the mysteries but there is no such expertise in Australia.

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REFERENCES Exon, N.F.; Bernadel, G.; Brown, J.; Cortese, A.; Findlay, C.; Hoffmann, K.; Howe, R.; Quilty, P. 2006a: The geology of the Mellish Rise region off northeast Australia: a key piece in a tectonic puzzle. Geoscience Australia Record 2006/08. 204 pp. Exon, N.F.; Hill, P.J.; Lafoy, Y.; Heine, C.; Bernadel, G. 2006b: Kenn Plateau off northeast Australia: a continental fragment in the southwest Pacific jigsaw. Australian Journal of Earth Sciences 53: 541-564. Hoffman, K.L.; Exon, N.F.; Quilty, P.G.; Findlay, C.S. 2008: Mellish Rise and adjacent deepwater plateaus off northeast Australia: new evidence for continental basement from Cenozoic micropalaeontology and sedimentary geology. In Blevin, J.E.; Bradshaw, B.E.; Uruski, C. ed. Eastern Australasian Basins Symposium III, Petroleum Exploration Society of Australia, Special Publication, Pp. 317-323. Pujalte, V.; Payros, A.; Apellaniz, E. 2009. Climate and biota of the Early Paleogene: recent advances and new perspectives. Geologica Acta 7: 1-9.

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ANALYSIS OF LATE PALEOCENE–EARLY EOCENE MICROPLANKTON FROM THE KHEU RIVER SECTION, WEST PRE-CAUCASUS E.P. Radionova1, G.N. Aleksandrova1, T.T. Gavtadze2, S.I. Stupin1 and I.E. Khokhlova3 1

2

Geological Institute, Russian Academy of Science, 119017, Pyzhevskii 7, Moscow: [email protected]; LEPL Aleksandr Djanelidze Institut of Geology, MAleksidse st.1/9, 0193 Tbilisi, Georgia; 3391c Oceanbeach Road, Tauranga, New Zealand: irina_khokhlova@ mail.ru

INTRODUTION The Kheu River section is an important Paleogene sequence which serves as the stratotype for the foraminiferal zonal scheme for central Pre-Caucasus. Krasheninnikov and Muzylev (1975) provided the first correlation between Paleogene foraminiferal and nannoplankton zonal schemes. In the upper Paleocene (base of Zone NP9b) at the Kheu River section is a 35 cm sapropel bed (further referred to as sapropelite A) that was first described by Muzylöv and Tabachnikova (1989). This lithological event marks the lower boundary of Zone NP9b in the vast CrimeaCaucasian and Middle Asia area (Muzylöv 1994) and, subsequently, accumulation of sapropelite A was shown to correspond with the PETM event (Gavrilov et al. 2000; Gavrilov and Shcherbinina 2004). We have studied microplankton across the Late Paleocene–Early Eocene boundary interval in the Kheu River section, near Nalchik, including nannofossils, foraminifera, siliceous plankton (i.e. diatoms, silicoflagellates, radiolarians) and dinoflagellate cysts (dinocysts). Analyses were conducted on the same sample suite, which provides correlation between the various microfossil groups. A primary aim of this study is analysis of microplankton assemblages during the PETM warming event. MATERIAL AND METHOD The lithologic composition of the Kheu River section is shown in Figure 1. The Nalchik Formation is subdivided into three units. The lower unit (I) is ca 20 m thick, and represented by greenish-grey soft marls and by light-green marls in the middle part. Foraminifera, nannoplankton, siliceous microplankton (diatoms) and sporadic dinocysts appear at the base of unit I. The second unit (II) is ca 6 m and composed of dark-grey carbonate-siliceous clays with indistinctive sapropel layers. The Extended Abstracts

base of unit II is marked by sapropelite A (at present, this part of the section is not exposed). In this unit, diatom (often covered with a nontransparent organic or pyrite compound) abundance increases, and radiolaria and dinocysts appear. The upper unit (III) is ca 20m thick and contains siliceous clays intercalated with dark slightly carbonate clays with gaize lenses, and layers containing numerous, mostly altered, radiolaria. RESULTS Nannoplankton In the Paleocene interval (samples 50 to 100) at the Kheu River section, we recognised six stratigraphic units (zones and subzones) that correspond to nannoplankton zones of the Crimea-Caucasian Region. The section is located near the boundary of a warm and moderate-cold climate belt, which resulted in a mixed assemblage of both warm- and coldwater species. This has allowed us to correlate nannoplankton assemblages from the Kheu River section with both high- and mid-latitude associations. The nannoplakton zonation of the Crimea-Caucasian area is primarily compiled from the standard schemes of Martini (1970) and Okada and Bukry (1980). It was possible to establish additional zonal subdivisions in the interval from the first appearance of Discoaster multiradiatus (end of the Paleocene) to the first appearance of D. lodoensis at the base of the lower Eocene (Fig.1). The nannoplankton zones established in the Kheu River section are almost identical to the zonation scheme of Aubry et al. 2002. Various studies have shown that the presence of the Rhomboaster-Discoaster araneus (RD) group in the Tethys region corresponds to the PETM interval (e.g. Aubry 2001; Aubry et al. 2007). In the Kheu River section, the FO of Discoaster araneus (Fig. 2) is found in sapropelite A (Stupin and Muzylöv 2001;

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Gavrilov and Shcherbinina 2004). Based on nannoplakton data, the PETM begins at the appearance of the RD group (and after the LO of Fasciculithus alanii) and terminates before the appearance of Discoaster mahmoudii. Hence, it may be predicted that the carbon isotope excursion occurs within the interval from sample 61 to 76 (i.e. Unit II). However, Tribrachiatus bramlettei first appears in sample 85 (Fig. 2), which suggests that the PETM may be expanded.

Foraminifera Planktonic foraminifera dominate (95%) the foraminiferal association in Unit I (Fig. 1). The moderate-warm Subbotina taxa are common, with warm-water Acarinina and Morozovella being rarer. The benthos is diverse (up to 45 species) and calcareous forms are most prevalent. The assemblage is typical of a “midtype” shelf fauna, with the presence of Anomalinoides welleri, A. acutus, Gavelinella

Figure 1 Sedimentary column of Nalchik Formation of the Kheu section with sampling points. Zonation and marker species on nannoplankton, foraminifera, diatoms, radiolarian and dinocysts. Unit II approximately corresponds to PETM interval.

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Extended Abstracts


danica, Cibicidoides succedens, Gyroidinoides subangulatus and Bulimina midwayensis, along with elements of a deepwater “Velasco-type” fauna (e.g. Nuttallides truempyi, Gavelinella beccariiformis). At the base of Unit II, close to sapropelite A, planktonic foraminifera become less abundant, and disappear completely ca 10 cm beneath the sapropel layer with pyrite-replaced shells of benthic taxa Bathysiphon sp. and Rhabdammina sp. present. In the lower and upper parts of sapropelite A (Fig. 2), planktonic foraminifera are dominant (99 %). In the lower part, Subbotina dominate but their abundance decreases upsection and warmwater taxa Acarinina and Morozovella become more abundant. The benthos is sparse, with the presence of calcareous agglutinated forms tolerant of low oxygen conditions (Kaiho 1991, 1992, 1994) such as Hippocrepinella sp., Haplophragmoides sp, Lenticulina sp., Bulimina sp. and Paradentalina sp.. In the middle part of the sapropel, the benthicplanktonic ratio reaches ca 60%. Many planktonic and benthic shells in the sapropel are pyrite-replaced and squeezed. Above sapropelite A (15-20 cm) the foraminifera association has recovered (Fig. 1). In the benthos, simple agglutinated foraminifera (e.g. Rhabdammina, Bathysiphon, Ammodiscus, Repmanina) are dominant and calcareous benthic forms become sparse. Planktonic foraminifera are less common and comprise mainly Subbotina forms, with fewer Acarinina and sparse Igorina and Morozovella. Unit III also contains agglutinated benthic foraminifera. From the base of Zone NP10a (sample 85), the structure of the planktonic and benthic assemblages is restored and the abundance and diversity of planktonic forms increases up-section. Subbotina is dominant and more Acarinina taxa (e.g. A. nitida, A. soldadoensis, A. coalingensis) are recorded. Diatoms and Silicoflagellates Diatoms and silicoflagellates appear at the base of Unit I (sample 50), and are sparse and represented by a pelagic assemblage (Fig. 1). The most abundant diatom genera are Paralia, Coscinodiscus, Trinacria, Stephanopyxis and Pyxidicula. Marker species Pyxidicula moelleri and Trinacria ventriculos are recorded in Unit I, which corresponds to the lower boundary of Extended Abstracts

nannoplankton Zone NP8. Silicoflagellates are represented by Dyctiocha elongata and D. precarentis. Provisionally, the lower to mid part of Unit I can be linked with the Trinacria ventriculosa Zone (Strelnikova 1992). The base of the Moissevia uralensis Zone is recognised in sample 57, within Unit I (Fig. 1). Diatoms become more frequent in Unit II and assemblage changes are noted – Paralia, Coscinodiscus and Craspedodiscus dominate (they are mostly pyrite-altered). Coscinodiscus denarius is abundant in the lower part (sample 62) of Unit II, which corresponds to the Hemiaulus proteus Zone (index species is found in sample 75). The first appearance of Coscinodiscus payeri and Pyxilla sp. also occurs in Unit II. Diatoms are rare in Unit III and the assemblage is represented by small shells of the genera Paralia, Trinacria and Stephanopyxis. The association is assigned to the Coscinodiscus payeri Zone. Radiolaria The assemblage of radiolaria species Petalospyris foveolata, Lychnocanium auxilla, Podocyrtis papalis, Heliodiscus sp., Thyrsocyrtis hirsuta tensa, Thyrsocyrtis cf. mongolfieri in sample 61, ca 1–1.5 m below the sapropel layer, is assigned to the upper Paleocene Petalospyris foveolata Zone (Boreal scheme of Kozlova 1990). This association resembles those of the same age in the Middle Volga and West Siberia, but contains more warm-water elements and the overall character points to a warm moderate (subtropical) climate. An abundant assemblage and good preservation suggests open ocean normal salinity conditions. Close to the sapropel layer radiolaria are rare and badly preserved, and above the sapropel layer (sample 75) the assemblage is mostly represented by Spongodiscidae, Prunoidea and small Spumellaria. This indicates a change in Paleogeographic environment after PETM towards shallower conditions with more restricted ocean connections and coarse accumulation process. Dinocysts In unit I, dinocysts are only found in sample57, in the upper part of nannoplankton Zone NP8 (1.5 m below the FO of D. multiradiatus). The

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sample contains an acme of Cyclonephelium sp. A, along with a single presence of Alisocysta margarita and Achomosphaera spp. In this study, dinocysts from unit II were examined from the mid to upper part of the "sapropel" deposit (samples 62–75; Fig. 1). According to N.I. Zaporozhets (pers. comm.), an Apectodinium acme (>70%) and the species Apectodinium augustum are recorded in the lower part of sapropelite A, at the base of Unit II and FO of P. plana (Fig. 2). Apectodiniumdominated assemblages have been shown to occur in the PETM from sites around the world (Crouch et al. 2001; Sluijs et al. 2007), which allows us to place Paleocene–Eocene boundary at the base of Unit II (Fig. 2). Above sapropelite A, the Apectodinium acme is not observed and Apectodinium augustum is found sporadically (sample 75). Overall, the dinocyst assemblage is similar to the upper Thanetian association of the Apectodinium hyperacanthum Zone (Powell 1992). A difference includes the distribution of Glaphyrocysta sp. B (= ex gr. GlaphyrocystaChiropteridium) to Unit II, and the RD nannoplankton complex differentiates this dinocyst assemblage from an upper Thanetian association, which allows us to limit its stratigraphic range to the PETM interval. The lower part of Unit III (samples 77–82) is characterized by dinocyst assemblages with low abundance and diversity. The base of Unit III is ca 7 m below the FO of Tribrachiatus bramlettei and contains the FO of Wetzeliella sp. 1, Tectatodinium pellitum and Cerebrocysta sp. (Fig. 1). In the middle part of unit III (samples 83–87) the Wetzeliella astra zone is recorded. This occurs ca 2.5 m below the FO of Tribrachiatus bramlettei and is defined by the FO of W. astra, Deflandrea truncata and Heteraulacacysta leptalea. The upper part of Unit III (samples 88–95) corresponds to the Wetzeliella meckelfeldensis zone, which is defined by the FO of W. meckelfeldensis (Powell 1992). Dinocyst assemblages have low diversity and abundance, and contain Achomosphaera alcicornu (prevails in sample 88), Achomosphaera spp., Spiniferites spp., Cyclonephelium sp. A, W. meckelfeldensis, Wetzeliella spp., ?Danea sp. and Deflandrea truncata. In the lower part of the W. meckelfeldensis Zone is the FO of Impagidinium crassimuratum, which 114

dominates in samples 94–95 along with D. truncata, Wetzeliella spp., Achomosphaera spp. and Spiniferites spp. At the Kheu River section, the lower boundary of the W. meckelfeldensis Zone coincides with the base of the Discoaster diastypus Zone (NP10d, Aubry et al. 2007) and is marked by a layer of rare siderite concretions. DISCUSSION The Kheu River section contains a continuous Late Paleocene–Early Eocene sequence of sediments and we have examined various microfossil groups and correlated the faunas to Atlantic and Western Europe zonal schemes. In the Late Thanetian, from nannoplankton Zones NP8–NP9a, sediments (Unit I) accumulated in an outer shelf–uppermost continental slope environment, in conditions of active surface water circulation, marine transgression and subtropical climate. Within carbonate and siliceous plankton groups this interval is characterized by maximum diversity, and an almost total absence of a dinocyst flora. The benthic foraminifera association is represented by a Midway-type assemblage, although there are also elements of Velasco-type assemblage. Sediments representing the PETM, according to nannoplankton, are marked by the stratigraphic range of the RD group and the lower part of the Romboaster cuspis (NP9b) Zone (Aubry 2001; Aubry et al. 2007). The PETM interval is also correlated to the lower part of the foraminiferal Acarinina nitida Zone, diatom Hemiaulus proteus Zone and dinocyst Apectodinium augustum Zone. The base of the PETM in the Kheu River section is marked by sapropelite A. In the bottom and top parts of the sapropel layer planktonic foraminifera dominate (up to 99%). Benthic foraminifera are sparse, with agglutinated forms Hippocrepinella sp., Haplophragmoides sp., and carbonaceous taxa Lenticulina sp., Bulimina sp. and Paradentalina sp. suggesting a tolerance of low oxygen content. Above sapropelite A agglutinated foraminifera dominate and the anotable regression and possible cooling of climate. The assemblage structure of planktonic groups changes, with the abundance

Extended Abstracts


Figure 2 Distribution of stratigraphic PETM interval markers in Sapropelite A (based on data from Stupin and Muzylöv 2001 and Gavrilov and Shcherbinina 2004), and in Unit II (based on data from this study). Arrows show levels of zonal nannoplankton species appearances. Extended Abstracts

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of carbonaceous plankton decreasing to 30%, increasing siliceous plankton (radiolaria) abundance and sparse diatoms represented by boreal species. From the base of Zone NP10a, the structure of planktonic and benthic foraminiferal assemblages is restored. In the upper part of the Acarinina nitida Zone (= Tribrachiatus contortus Zone, NP10b–d) warm water Acarinina species appear. In the dinocyst assemblage, from Zone NP9b to Zone NP10a, there is a change from the upper part of the Apectodinium hypercanthum Zone to the Wetzeliella meckelfeldensis Zone. Species that belong to the Wetzeliellaceae family (Apectodinium, Wetzeliella) demonstrate rapid evolution — Wetzeliella sp. 1 is possibly a new transitional species in the Apectodinium–Wetzeliella line. REFERENCES Aubry, M.-P. 2001: Provincialism in the photic zone during the LPTM. In Ash, A.; Wing. S. ed. Climate and Biota of the Early Paleogene, International meeting, Powell, Abstract Volume. Pp. 6. Aubry, M.-P.; Working Group on the Paleocene– Eocene boundary 2002: The Paleocene–Eocene boundary global stratotype section and point (GSSP): criteria for characterization and correlation. Tertiary Reserch 21: 57-70 Aubry, M.-P.; Ouda, K.; Dupuis, C.; Berggren, W.A.; Van Couvering, J; Working Group on the Paleocene–Eocene boundary 2007: The global stratotype section and point (GSSP) for the base the Eocene Series in the Dababiya section (Egypt). Episodes 30: 271-286 Crouch, E.M.; Heilmann-Clausen, C.; Brinkhuis, H.; Morgans, H.; Rogers, K.; Egger, H.; Schmitz, B. 2001: Global dinoflagellate event associated with the Late Paleocene Thermal Maximum. Geology 29: 315–318. Gavrilov, Y.O.; Shcherbinina, E.A.; Muzylöv, N. 2000: A Paleogene sequence in central North Caucasus: a response to paleoenvironmental changes. GFF 122: 51-53. Gavrilov, Y.O.; Scherbinina, E.A. 2004: Global biosphere event at the Paleocene–Eocene Boundary. In Gavrilov, Y.O.; Khutorskoy, M.D.; ed. Modern problems of Geology. “Nauka” Mosow. Transactions of the Geological Institute. Vol. 565. Pp. 493–531. Kaiho, K. 1991: Global changes of Paleogene aerobic/anaerobic benthic foraminifera and

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deep-sea circulation. Palaegeography, Paleoclimatology, Paleoecology 83: 65-85. Kaiho, K. 1992: Eocene to Quaternary bentic foraminifers and paleobathymetry of the IzuBonin Area, Leg 125 and 126. In Taylor, B.; Fujioka, K.; Janecek, T.; Lanngmuir, C. ed. Proceedings of the Ocean Drilling Program, Scientific Results 126. Pp. 285-310. Kozlova, G.E. 1999: Paleogene Boreal Radiolarians from the Russia. Ministry of Natural Resourses of Russian Federation. All-Russia Petroleum Research Exploration Institute (VNIGRI). 323 p. (in Russian) Krasheninnikov, V.A.; Muzylev, N. G. 1975: Relationship of zonal schemes by planktonic foraminifera and nannofossils in the Paleogene sections of Northern Caucasus. Voprosy Micropaleontologii 18: 212-224. (in Russian) Martini, E. 1971: Standard Tertiary and Quaternary calcareous nannoplankton zonation. In Farinacci, A. ed. Proceedings of the 2nd Planktonic Conference, Roma. Tecnoscienza Roma. Pp. 739-785. Muzylöv, N.G; Benyamovskii, V.N.; Tabachnikova, I.P. 1989: Sapropel interlayers in Lower Paleocene deposits of the Southern of Soviet Union. Izvestia Akademii Nauk SSSR, Seria Geologitcheskaya 11: 117-119. (in Russian) Muzylöv, N.G. 1994: Anoxic events of Paleocene– Eocene. In Rozanov A.; Semikhatov M.A. ed. Ecosystem Restructures and the Evolution of Biosphere. ”Nedra” Moscow, Issue 1. Pp. 160166. (in Russian) Okada, H.; Bukry, D. 1980: Supplementary modification and introduction of code numbers to the low-latitude coccolith biostratigraphic zonation (Bukry, 1973, 1975). Marine Micropaleontology 5: 321-325. Powell, A.J. 1992: Dinoflagellates cysts of the Tertiary system: A stratigraphic index of dinoflagellates cysts. British Micropaleontological Society Publication. Pp. 152-251. Strelnikova, N.I. 1992: Paleogene Diatoms Algae. St-Petersburg State University Publication. 311p. (in Russian) Stupin, S.I.: Muzylöv, N.G. 2001: The Late Paleocene ecologic crisis in epicontinental basins of the Eastern Peri-Tethys: Microbiota and accumulation conditions of sapropelitic bed. Stratigraphy, Geological Correlation 9: 87-93. (in Russian)

Extended Abstracts


NEW ZEALAND PALEOGENE VEGETATION AND CLIMATE J.I. Raine, E.M. Kennedy and E.M. Crouch GNS Science, P.O. Box 30-368, Lower Hutt, New Zealand: [email protected]

New Zealand has a near continuous sedimentary record through the Cenozoic. A Paleogene pollen record is derived from both marine and terrestrial sediments and there are many Paleogene plant macrofossil localities, few as yet described in detail. Current key research questions include: does NZ vegetation show abrupt change at the Paleocene–Eocene Thermal Maximum (PETM); does plant fossilderived climate data corroborate tropical sea surface temperature results from marine faunas and geochemical proxies for the Early Eocene (Hollis et al. 2009); what happened to the flora during Middle and Late Eocene cooling events and in the Late Oligocene when most land was submerged (Landis et al. 2008)? Following ca 20 million years of rifting and crustal thinning during the Late Cretaceous, the New Zealand continental mass (Zealandia) split from Australia-Antarctica at about 85 Ma. A phase of passive submergence ensued during Tasman Sea opening 85–55 Ma, with extensive marginal coastal plain and shelfal marine deposition. During the Middle to Late Eocene a shift in plate boundaries resulted in seafloor spreading south of New Zealand and subduction to the north, with local compression and uplift in Zealandia but generally renewed submergence and marine transgression. Inception of the current convergent plate boundary in the Early Miocene was accompanied by emergence and widespread terrestrial sedimentation, especially in the South Island, although Zealandia today remains 95% submerged. During Tasman Sea and Southern Ocean spreading the present New Zealand landmass gradually moved northwards to mid latitudes: 60–70°S at 65 Ma, 52–62°S at 55 Ma, 45–55°S at 40 Ma, and 42–52°S at 25 Ma (Yan and Kroenke 1993). From the Late Cretaceous the flora began to diverge from that of Australia, although continuing migration acted to reduce dissimilarity. For example, pollen of Nothofagus (southern beech) subgenera Nothofagus and Brassospora appeared in Extended Abstracts

both Australia and New Zealand during the Eocene, either by trans-oceanic dispersal or via a land bridge to Australia that may have persisted through the northern Lord Howe Rise. General trends in vegetation change are reflected in the palynological zonation of Raine (1984, 2004; Morgans et al. 2004), with three superzones characterised by successive abundance of pollen of podocarps (Phyllocladidites mawsonii Assemblage, Late Cretaceous to basal Eocene), Casuarinaceae (Myricipites harrisii Assemblage, Early to Late Eocene) and Brassospora (Nothofagidites matauraensis Assemblage, Late Eocene to Oligocene). Pollen diagrams from Tui-1 (Figs. 1 and 2) and Huntly Coalfield (Fig. 3) illustrate these changes. The Cretaceous–Cenozoic boundary event caused reduction in pollen diversity, particularly in angiosperms, although after the

Figure 1

Localities mentioned in text.

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Figure 2 Summary pollen diagram for Tui-1 petroleum exploration well, southern Taranaki Basin (based on Pollock et al. 2003). The relative abundance of dinoflagellate cysts also shows the degree of marine influence in these lower coastal plain to inshore marine sediments, notably the basal Eocene flooding also reflected in high levels of Nypa pollen. Many samples are drill cuttings, so there is some carry-down of miospores to greater depths. For New Zealand stage abbreviations and pollen zones, see Morgans et al. (2004).

initial revegetation of the land by pteridophytes and gymnosperms some opportunistic angiosperms such as Triorites minor became more abundant (Vajda et al. 2001, 2003; Vajda and Raine 2003). The general character of the Paleocene palynoflora (PM3a zone) remained similar to that of the Late Cretaceous, with abundant conifers (mainly Podocarpaceae and Araucariaceae); among the angiosperms, Proteaceae are a relatively common component with known modern affinity. The palynoflora is fairly uniform through the Paleocene, but increasing pteridophyte abundance seen in the Tui-1 section may reflect gradually increasing temperatures. Early Paleocene macrofossil localities include several at Pakawau (NW Nelson) and Greymouth Coalfield (Westland); other Paleocene localities include the Canterbury region localities Kakahu and Mt Somers Coal Mine (Pole 1997, 1998). In contrast to the palynofloras, leaf assemblages tend to be angiosperm-dominated, with species of Proteaceae and Lauraceae prominent, but they too suggest lower diversity relative to the Late 118

Cretaceous and slightly cooler, temperate climate. Analyses of leaf morphology (LMA and CLAMP techniques) from three Early Paleocene localities (Kennedy 2003) indicate ca 6 to 12°C mean annual temperature, with growing season length (months with mean temperature ≥10°C) of 6.4 to 7.6 months and mean growing season precipitation of over 2000 mm. Warmer climate indicators begin to appear in the pollen record at or just prior to the PETM – e.g. Cupanieidites (Sapindaceae: Cupania), Malvacipollis (Euphorbiaceae: Austrobuxus) and Spinozonocolpites (Arecaceae: Nypa). Earliest Eocene spore-pollen assemblages (PM3b zone) are of mixed Paleocene–Eocene character, with continued abundance of conifer pollen, for example at the Tawanui section (Crouch and Visscher 2003). High sea level at the time of the P–E boundary is reflected in increased marine influence in several paralic sequences, e.g. Kumara-2 core (Sluijs et al. 2008) and Mt Somers sand quarry (Raine and Wilson 1988; Kennedy et al. 2006) (see also Fig. 2). Known macrofossil assemblages Extended Abstracts


across the P–E transition are scarce. The Otaio River leaf flora from coal measures in South Canterbury is possibly closest to the PETM: associated Apectodinium suggests the leaf flora occurs within one of the Apectodinium dinocyst acmes that occurred during and after the PETM (Crouch and Brinkhuis 2005). Preliminary Leaf Margin Analysis suggests a warm climate. Moreover, based on present day ecological tolerances (Fechner 1988), the mangrove palm Nypa cannot survive in temperatures of less than 20°C. Later Early to Middle Eocene palynofloras (MH1 zone) are dominated by pollen of Casuarina, with Proteaceae also prominent and diverse. Nypa pollen is frequent in estuarine sediments, and Anacolosidites (Olacaceae: Anacolosa), and Schizocolpus (Didymelaceae) are additional tropical elements. In reviewing palynological evidence for Early and Middle Eocene climate, Pocknall (1990a) suggested coastal temperatures of 20–24°C (mesothermal/megathermal interzone of Nix, 1982), and recognised the presence of a mangrove association, local wetlands with common Liliaceae and Gleicheniaceae, coastal/lowland sclerophyll scrub dominated by Casuarinaceae and Proteaceae, and a hinterland association characterised by rainforest elements. The Livingstone macrofossil flora from North Otago may correlate to the MH1 zone: Pole (1994) described 13 angiosperm leaf taxa (Myrtaceae, Proteaceae, Sterculiaceae, Elaeocarpaceae, possibly Lauraceae, plus several unattributed forms) and two conifers, and reached a similar conclusion to that of Pocknall in inferring a warm climate with seasonal water-stress. Cooling in the late Middle Eocene (New Zealand Porangan Stage) is heralded by the reappearance of common Nothofagus pollen, rare since the Late Cretaceous. Nothofagidites flemingii (Nothofagus subgenus Nothofagus) characterises zone MH2, but representatives of subgenera Fuscospora and Brassospora are also present. A short-lived drop in Nothofagus pollen abundance during the early Late Eocene (zone MH3) may represent a warmer interlude. This zone is known from northern and western New Zealand but appears to be absent in southern New Zealand (Southland and Fiordland). Leaf fossils from the MH2 and MH3 zones occur in the widespread Brunner

Extended Abstracts

Coal Measures of Westland, but have been little studied. Towards the close of the Eocene there was an abrupt change to Brassospora-dominated palynofloras (Nothofagidites matauraensis Assemblage), which also characterise the Oligocene. This change is not precisely dated, but appears to occur within the NZ Runangan Stage, therefore slightly predates the Eocene– Oligocene boundary (Morgans et al. 2004). It is associated with progressive marine transgression, but because of its abruptness and independence of depositional facies must have a regional climatic origin. Pocknall (1989) argued that a modern analogue to the latest Eocene vegetation may be the cool-temperate rainforests of the New Guinea highlands, where rainfall is at least 1500 mm per annum, mean annual temperature 13–18°C, and there is minor seasonal variation. Leaf floras are known from widespread Late Eocene to Early Oligocene coal measures, e.g. in Waikato (Unger 1864, Penseler 1930), but again have been little studied in recent decades. From the Late Eocene Beaumont Coal Measures at Pikopiko, West Southland, Lee et al. (2003, 2009) have recorded leaves of Nothofagus, Metrosideros, Lauraceae, Calamus (rattan palm), and a liane (Smilax or Ripogonum). Abundant and diverse ferns and epiphyllous fungi support interpretation of a perhumid but warm temperate climate. The Eocene–Oligocene boundary itself has not been recognised in New Zealand terrestrial sequences. In the Huntly Coalfield at about this level a zone of relatively abundant Myrtaceae pollen (NM1) is succeeded by an acme of Araucariaceae (lower zone NM2, Fig. 3). A similar succession is seen in Southland at the possibly correlative transition from Beaumont Coal Measures to the overlying lacustrine Orauea Mudstone (Pocknall and Turnbull 1989), but the climatic significance of these changes is unknown. Oligocene macrofossil floras are at present undocumented although they could occur within the upper Waikato Coal Measures of the Raglan area (Pocknall 1985) and lower Gore Lignite Measures of Southland (Pocknall 1990b; Pocknall and Mildenhall 1984). Oligocene palynofloras have been documented from several regions, however precise dating remains problematic. Within the Brassospora-dominated Late

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Su b O tro th p er ic an al a gi ng Sp o s io or pe sp es rm er s ms

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Figure 3 Summary pollen diagram based on core from Huntly Coalfield drillhole 7153, Waikato (data from Pocknall 1991; zones after Raine 2004). Strata in the section overlie coal measures of zone MH3 (NZ Kaiatan– early Runangan stages) elsewhere in the coalfield; marine strata at the top of the section (Whaingaroa Siltstone) are of late Early to early Late Oligocene age (NZ Upper Whaingaroan Stage).

Oligocene palynofloras from Southland Nothofagidites cranwelliae succeeded N. matauraensis as the most common pollen type. Forests appear to have remained dominated by Nothofagus and gymnosperms, but increased floristic diversity in the Late Oligocene may reflect slightly increased temperatures. It can be seen that there is a good deal of qualitative floral data to support warming climate at the Paleocene–Eocene boundary and during the Early Eocene, and later fluctuations including cooling in the Middle Eocene and near the Eocene–Oligocene boundary. In general, the observations support temperature results from marine faunas and geochemical proxies but there remain questions of precise correlation, discrimination of local from regional effects, and derivation of precise estimations of climate from fossil plant occurrences. Nevertheless, there is plenty here to attract further study.

from eastern New Zealand: palynological approach. Micropaleontology 56: 138-160.

a

marine Marine

Crouch, E.M.; Visscher, H. 2003: Terrestrial vegetation record across the initial Eocene thermal maximum at the Tawanui marine section, New Zealand. Special paper, Geological Society of America 369: 351-363. Fechner, G.G., 1988: Selected palynomorphs from the Lower to Middle Eocene of the South Atlas Border Zone (Morocco) and their environmental significance. Palaeogeography, Palaeoclimatology, Palaeoecology 65: 73-79. Hollis, C.J.; Handley, L.; Crouch, E.M.; Morgans, H.E.G.; Baker, J.A.; Creech, J.; Collins, K.S.; Gibbs, S.J.; Huber, M.; Schouten, S.; Zachos, J.C.; Pancost, R.D. 2009: Tropical sea temperatures in the high-latitude South Pacific during the Eocene. Geology 37: 99-102.

REFERENCES

Kennedy, E.M. 2003: Late Cretaceous and Paleocene terrestrial climates of New Zealand: leaf fossil evidence from South Island assemblages. New Zealand Journal of Geology and Geophysics 46: 295-306.

Crouch, E.M.; Brinkhuis, H. 2005: Environmental change across the Paleocene–Eocene transition

Kennedy, E.M.; Crouch, E.M.; Raine, J.I.; Pancost, R.D.; Handley, L. 2006: Paleocene-Eocene

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transition in South Island terrestrial to marginal marine sections. Geological Society of New Zealand, New Zealand Geophysical Society Joint Conference, Palmerston North, 2006, programme and abstracts. Geological Society of New Zealand Miscellaneous Publication 122A: 36-37. Landis, C.A.; Campbell, H.J.; Begg, J.G.; Mildenhall, D.C.; Paterson, A.M.; Trewick, S.A. 2008: The Waipounamu erosion surface: questioning the antiquity of the New Zealand land surface and terrestrial fauna and flora. Geological Magazine 145: 173-197. Lee, D.; Lindqvist, J.; Douglas, B.; Bannister, J.; Cieraad, E. 2003: Paleobotany and sedimentology of Late Cretaceous-Miocene nonmarine sequences in Otago and Southland: field trip 9, Geological Society of New Zealand annual conference 2003. In Cox, S.C.; Smith Lyttle, B. ed. Geological Society of New Zealand Inc 2003 annual conference, University of Otago, Dunedin: field trip guides. Geological Society of New Zealand Miscellaneous Publication 116B: 48 p. Lee, D.E.; Lindqvist, J.K.; Bannister, J.M.; Cieraad, E.; Raine, J.I.; Kennedy, E.M.; Conran J.C. 2009: Late Eocene climate of southern New Zealand: insights from the in situ Pikopiko Fossil Forest. In Strong, C.P.; Crouch, E.M.; Hollis, C.J. ed. Climatic and Biotic Events of the Paleogene conference, Wellington, New Zealand, January 12th-15th, 2009, Programme and abstracts. GNS Science Miscellaneous Series 16. Pp. 112. Morgans, H.E.G.; Beu, A.G.; Cooper, R.A.; Crouch, E.M.; Hollis, C.J.; Jones, C.M.; Raine, J.I.; Wilson, G.J.; Wilson, G.S. 2004: Paleogene. Chapter 11 In Cooper, R.A. ed. The New Zealand Geological Timescale. Institute of Geological and Nuclear Sciences Monograph 22. Pp. 124-161. Nix, H. 1982: Environmental determinants of biogeography and evolution in Terra Australis. In Barker, W.; Greenslade, P. ed. Evolution of the flora and fauna of arid Australia. Adelaide, Peacock Publ. Pp. 47-66. Penseler, W.H.A. 1930: Fossil leaves from the Waikato district, with a description of the coal measure series. Transactions of the New Zealand Institute 61: 452-477. Pocknall, D.T. 1982: Palynology of Late Oligocene Pomahaka estuarine bed sediments, Waikoikoi, Southland, New Zealand. New Zealand Journal of Botany 20: 263-287 Pocknall, D.T. 1985: Palynology of Waikato Coal Measures (Late Eocene–Late Oligocene) from

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the Raglan area, North Island, New Zealand. New Zealand Journal of Geology and Geophysics 28: 329-349. Pocknall, D.T. 1989: Late Eocene to Early Miocene vegetation and climate history of New Zealand. Journal of the Royal Society of New Zealand 19: 1-18. Pocknall, D.T. 1990a: Palynological evidence for the early to middle Eocene vegetation and climate history of New Zealand. Review of Palaeobotany and Palynology 65: 57-69. Pocknall, D.T. 1990b: Palynology. In Isaac, M.J.; Lindqvist, J.K.; Pocknall, D.T. ed. Geology and lignite resources of the East Southland Group, New Zealand. New Zealand Geological Survey Bulletin 101: 141-152. Pocknall, D.T. 1991: Palynostratigraphy of the Te Kuiti Group (Late Eocene–Oligocene), Waikato Basin, New Zealand. New Zealand Journal of Geology and Geophysics 34: 407417. Pocknall, D.T.; Mildenhall, D.C. 1984: Late Oligocene-Early Miocene spores and pollen from Southland, New Zealand. New Zealand Geological Survey Paleontological Bulletin 51. 64 p. Pocknall, D.T.; Turnbull, I.M. 1989: Paleoenvironmental and stratigraphic significance of palynomorphs from Upper Eocene (Kaiatan) Beaumont Coal Measures and Orauea Mudstone, Waiau Basin, western Southland, New Zealand. New Zealand Journal of Geology and Geophysics 32: 371-378. Pole, M.S. 1994: An Eocene macroflora from the Taratu Formation at Livingstone, North Otago, New Zealand. Australian Journal of Botany 42: 341-367. Pole, M. 1997: Paleocene plant microfossils from Kakahu, south Canterbury, New Zealand. Journal of the Royal Society of New Zealand 27: 371-400. Pole, M. 1998: Paleocene gymnosperms from Mount Somers, New Zealand. Journal of the Royal Society of New Zealand 28: 375-403. Pollock R.M.; Strong, C.P.; Raine J.I.; Browne G.H. 2003: Biostratigraphy and paleoenvironmental interpretation of Tui-1 well, integrating a sedimentary and depositional interpretation of the Kapuni Group, Taranaki Basin, New Zealand. Institute of Geological and Nuclear Sciences Client Report 2003/85. 39 p. + appendix, encl. Raine, J.I. 1984: Outline of a palynological zonation of Cretaceous to Paleogene terrestrial sediments in West Coast Region, South Island,

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New Zealand. Report, New Zealand Geological Survey 109. 82 p. Raine, J.I. 2004: New Eocene and Oligocene miospore zones and subzones. Appendix to Chapter 11 In Cooper, R.A. ed. The New Zealand Geological Timescale. Institute of Geological and Nuclear Sciences monograph 22. Pp. 162-163. Raine, J.I.; Wilson, G.J. 1988: Palynology of the Mt Somers (South Island, New Zealand) early Cenozoic sequence (Note). New Zealand Journal of Geology and Geophysics 31: 385390. Sluijs, A.; Brinkhuis, H.; Crouch, E.; John, C.; Handley, L.; Munsterman, D.; Bohaty, S.; Zachos, J.; Reichart, G-J.; Schouten, S.; Pancost, R.; Damste, J.; Welters, N.; Lotter, A.; Dickens, G. 2008: Eustatic variations during the Palaeocene–Eocene greenhouse world. Paleoceanography 23: PA4216, doi:10.1029/2008PA001615. Unger, F. 1964: Fossile Pflanzenreste aus NeuSeeland. Reise d'Österreichischen Fregatte Novara um die Erde 1857-1859. Geologische Theil 1: 1-13.

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Vajda, V.; Raine, J.I. 2003: Pollen and spores in marine Cretaceous/Tertiary boundary sediments at mid-Waipara River, North Canterbury, New Zealand. New Zealand Journal of Geology and Geophysics 46: 255273. Vajda, V.; Raine, J.I.; Hollis, C.J. 2001: Indication of global deforestation at the CretaceousTertiary boundary by New Zealand fern spike. Science 294: 1700-1702. Vajda, V.; Raine, J.I.; Hollis, C.J.; Strong, C.P. 2003: Global effects of the Chicxulub impact on terrestrial vegetation - review of the palynological record from New Zealand Cretaceous/Tertiary boundary. In Dypvik, H.; Burchell, M.; Claeys, P. ed. Cratering in Marine Environments and on Ice. Impact Studies. Berlin, Springer Verlag. Pp. 57-74. Yan, C.Y.; Kroenke, L.W. 1993: A plate tectonic reconstruction of the southwest Pacific, 0–100 Ma. In Berger, W.H., Kroenke, L.W., Mayer, L.A., et al., ed. Proceedings of the Ocean Drilling Programme, Scientific Results 130. College Station, Texas, Ocean Drilling Program. Pp. 697–709.

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HIGH-RESOLUTION PALEOCENE STUDY OF NANNOFOSSIL AND DINOCYST ASSEMBLAGES FROM THE EASTERN CAUCASUS, SOUTHERN RUSSIA E.A. Shcherbinina and G.N. Aleksandrova Geological Institute of Russian Academy of Sciences, Pyzhevsky 7, Moscow 119017, Russia: [email protected]

INTRODUCTION The extensive epeiric basin of the northeastern Peri-Tethys contains a wide range of Paleocene sediments. The shallower northern part of the basin contains mainly siliciclastic sediments with good organic-walled microplankton assemblages, while the southern deepest area is dominated by a carbonate facies rich in calcareous microplankton. Sections that contain both dinocysts (DN) and nannofossils (NF) are rare, which has hindered the correlation of zonal subdivisions between these two biostratigraphically important microfossil groups. Typically, Danian sediments in the southern part of the NE Peri-Tethys (Crimea, Caucasus and Central Asia) are primarily limestones, with a cyclic appearance. The Danian– Selandian transition shows a drastic change to marlstone, calcareous clay (various formations) or non-calcareous clay (Goriachiy Kluch Formation), along with hiatuses of varying magnitude in different regions of the basin. Poor DN and sparse NF assemblages were recovered from the Goriachiy Kluch Formation, which provided recognition of a coarse DN zonation and tentative correlation with NF zones (Andreeva-Grigorovich 1992). During our field trip in the eastern Caucasus, we recovered one of the most complete Paleocene sections in the northern Urma Plateau, central Dagestan (Fig.1), which contains good assemblages of both dinocysts and nannofossils (Fig.2) The upper Danian part of Straus Ferma section is built up of massive layered limestone, in places disturbed by submarine slumping, with sporadic intercalations of calcareous shales (thickness studied ~30 m). The uppermost 5 m of this limestone succession is characterized by an

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aleuritic admixture. The Danian–Selandian transition corresponds to a hardground surface and a gradual change to more clayey sedimentation. Selandian–lower Thanetian deposits are composed of a monotonous soft greenish marlstone (thickness exposed ~35 m). RESULTS Nannofossils Nannofossils vary greatly in abundance and preservation. They are sparse and poorly preserved in the Danian, but become more abundant and diverse toward the top of Thanetian. The oldest NF assemblage examined (samples 140-141; Fig. 1) contains rare Prinsius martini, P. dimorphosus, Coccolithus pelagicus, C.robustus, Neochiastozygus modestus, Placozygus sigmoides, Chiasmolithus danicus, Cruciplacolithus primus, C. subrotundus, Toweius sp., consistent Braarudoaphaera bigelowii, B. discula, micrantoliths and calcareous dinocysts Thoracosphaera spp., and is assigned to Subzone NTp7A of Varol (1989). Marker species for the upper Danian Zone NP4 of Martini (1971) are not recognised in this interval. Upsection (13 m of section omitted due to intensive submarine slumping) the first occurrence (FO) of Chiasmolithus edentulus is recorded, which indicates the base of Subzone NTp7B. The first radiation of fasciculiths, which has been recorded at this time in many regions (e.g. Guasti et al. 2006; Steurbaut and Sztrakos 2008; Bernaola et al. 2009; Spronge et al. 2009), is difficult to detect in the Straus Ferma section. It may possibly occur within the 13 m of section not examined, because only a few specimens of Fasciculithus sp. were found in sample 142.

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Figure 1 The Paleocene interval of the Straus Ferma section, central Dagestan, eastern Caucasus, showing the lithology, modern and paleogeographic localities, and correlation of NF and DN events. 1 – zonation of Martini (1971); 2 – zonation of Varol (1989).

The FO of common Sphenolithus primus, which marks the base of Zone NTp8, is recorded in sample 143 (Fig. 1), where it corresponds to the FOs of Bomolithus elegans, Toweius sp., a large Prinsius martini and a significant increase in the abundance and 124

diversity of NF. Such a notable change in the NF assemblage may be caused by a hiatus or, more likely, by coarse sampling. From this level upwards, redeposited Cretaceous species are present to the top of the Danian. The FO of Neochiastozygus perfectus s.s. is recorded in Extended Abstracts


the middle of Subzone NTp8A (Fig. 1). The uppermost 5 m of Danian (silty limestone) is characterised by a decrease in NF abundance, the last occurrence of Braarudospaheara and Micrantolithus pentaliths, and domination of Toweius spp. over Coccolithus spp.. The coeval FOs of Fasciculithus ulii and F. janii (base of NTp8b and NTp8c, uppermost

Danian, respectively), along with the FO of F.tympaniformis (base of NP5, Selandian), just above the hardground surface suggests that a short hiatus occurs at the Danian–Selandian boundary, corresponding to zones NTp8bNTp8c. The Selandian–lower Thanetian interval is

Figure 2 Calcareous nannofossils and dinocysts from Paleocene of Straus Ferma section. Photos 1–17; Nannofossils, cross-polarized light, scale bar is 5 mkm in all images: 1 – Fasciculithus tympaniformis, sample 148/5; 2 – large Prinsius martinii, sample 143; 3 – Chiasmolithus edentulus, sample 148/4; 4 – Bomolithus elegans, sample 144; 5 – Heliolithus kleinpellii, sample 163; 6 – Fasciculithus involutus, sample 151; 7 – Toweius tovae, sample 152; 8 – Micrantolithus pinguis, sample 140; 9 – Neochiastozygus perfectus, sample 148; 10 – Hornibrookina sp., sample 162; 11 – Coccolithus robustus, sample 147; 12 – Fasciculithus ulii, sample 151; 13 – Fasciculithus janii, sample 148/6; 14 – Fasciculithus thomasii, sample 152; 15 – small Prinsius martinii, sample 142; 16 – Toweius pertusus, sample 150; 17 – Neochiastozygus saepes, sample 152. Photos 18–27; Dinocysts, scale bar is 20 mkm in all images: 18 – Spinidinium densispinatum, sample 152; 19 – Cladopyxidium saeptum, sample 159; 20 – Xenicodinium lubricum, sample 145; 21 – Impagidinium velorum, sample 153; 22 – Hafniasphaera cryptovesiculata, sample 147; 23 – Cerebrocysta sp. 1, sample 150; 24 – Hystrichosphaeridium sp. 1 Heilmann-Clausen, 1985, sample 144; 25 – Florentinia ferox, sample 155; 26 – Tectatodinium rugulatum, sample 146; 27 –, sample 143. Extended Abstracts

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characterized by abundant NF assemblages and a rapid diversification of Toweius, Fasciculithus and Neochiastozigus spp.. The FO of Heliolitus kleinpelli marks the base of NP6 zone, at 25.5 m above Danian–Selandian boundary. Dinocysts Dinocysts are not present in the lowermost part of interval studied (samples 140-142), but a diverse assemblage (~40 taxa) is recorded in sample 143. From this level to the top of Danian, the DN assemblage becomes more diverse and a series of FO datums are recognised. In the Selandian–lower Thanetian, the diversity of DN assemblage decreases significantly and the most stratigraphically important events in this interval are primarily LO datums. On the basis of the recorded DN event succession, a regional zonal subdivision is suggested. The oldest DN assemblage found in the Straus Ferma section appears to lie within the Alisocysta reticulata Zone, the top of which is defined by the LOs of A. reticulata and Xenicodinium lubricum. This assemblage (samples 143-146) is dominated by Spiniferites spp., Operculodinium (O. centrocarpum O. microtriania), Cordosphaeridium (C. inodes C. exilimurum) and Hystrichosphaeridium (H. tubiferum, Hystrichosphaeridium sp. 1 sensu Heilmann-Clausen 1985). Also present are Fibradinium anettoprense, Cerodinium diebelii, Fibrocysta ovalis, F. axiale, Danea mutabilis, Hystrichostrogylon coninckii. The top of the overlying Hafniasphaera cryptovesiculata Zone corresponds to LO of H. cryptovesiculata, which corresponds to the Danian–Selandian boundary. In the middle of this zone (samples 147-148), Hafniasphaera spp. becomes dominant. A dramatic change in the DN assemblage occurs at the base of the Selandian, with poor and low diversity (5–15 species) assemblages recorded. An abundant influx of redeposited Cretaceous spores and pollens is noted, while reworked DN are very rare. This feature is most likely due to enhanced erosion of on-land and shore-land Cretaceous deposits and, possibly, a switch of source area. The Cladopyxidium saepum Zone ranges from the LO of H. cryptovesiculata to the LO of C.

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saeptum. Species of Impagidinium (I. velorum, I. elegans, Impagidinium sp. 1 sensu Heilmann-Clausen, 1985), Spiniferites spp., Batiacasphaera spp., Cerebrocysta sp. 1 and Phthanoperidinium crenulatum are dominant in this interval (samples 149-162). The Cerodinium speciosum Zone spans from the LO of C. saeptum to the FO of Alisocysta margarita. In this interval, Impagidinium spp. and Phthanoperidinium crenulatum decrease in abundance, while Cerodinium speciosum increases in abundance up to sample 166 where its LO is recorded. The Alisocysta margarita Zone (lower boundary defined as the FO of A. margarita) is characterized by a sparse assemblage, which includes Spiniferites and Batiacasphaera spp. CONCLUSION The co-occurrence of diverse NF and DN assemblages in the Straus Ferma section has allowed us to directly correlate the zonal subdivisions of these two microfossil groups. The Alisocysta reticulata DN zone corresponds to the lower part of NF Subzone NTp8A (up to the FO of Neochiastozygus perfectus). The Hafniasphaera cryptovesiculata DN zone spans the upper part of NF Zone NTp8A. The Cladopyxidium saeptum DN zone covers the whole interval of NF Zone NP5, and Cerodinium speciosum DN zone corresponds to lower part of NF Zone NP6. ACKNOWLEDGEMENTS RFBR Project 4185.2008.5.

no.

09-05-00872;

SS-

REFERENCES Andreeva-Grigirovich, A.S. 1991: Phytoplankton (Dinocyst and Nannoplankton) Zonal Stratigraphy of the Paleogene of Southern USSR. Doctoral Dissertation in Geology and Mineralogy, Kiev University, SSSR). Bernaola, G.; Martin-Rubio, M.; Baceta J.I. 2009: New high resolution calcareous nannofossil analysis across the Danian/Selandian transition at the Zumaia section: comparison with South Tethys and Danish sections. Geologica Acta 7: 79-92. Guasti, E.; Speijer, R.P.; Brinkhuis, H.; Smit, J.; Steurbaut, E. 2006: Paleoenvironmental change at the Danian–Selandian transition in Tunisia: Foraminifera, organic-walled dinoflagellate Extended Abstracts


cyst and calcareous nannofossil records. Marine Micropaleontology 59: 210-229.

foraminifera Igorina albeari. Geologica Acta 7: 63-77.

Martini, E. 1971: Standard Tertiary and Quaternary calcareous nannoplankton zonation. In Farinacci, A. ed. Proceedings of the Second Planktonic Conference, Rome, (1970), Edizioni Tecnoscienza, 2. Pp. 739-785.

Steurbaut, E.; Sztrakos, K. 2008: Danian/Selandian boundary criteria and North Sea Basin–Tethys correlations based on calcareous nannofossil and foraminiferal trends in SW France. Marine Micropaleontology 67: 1–29.

Sprong, J.; Speijer, R.P.; Steurbaut, E. 2009: Biostratigraphy of the Danian/Selandian transition in the southern Tethys. Special reference to the Lowest Occurrence of planktic

Varol O. 1989: Palaeocene calcareous nannofossil biostratigraphy. In Crux, J.A.; van Heck, S.E. ed. Nannofossils and their Applications. British Micropalaeontological Society 12. Pp. 267310.

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EFFECT OF OCEAN GATEWAY CHANGES UNDER GREENHOUSE WARMTH. Willem P. Sijp Climate Change Research Centre, University of New South Wales, Sydney, Australia: [email protected]

The role of tectonic Southern Ocean gateway changes in driving Antarctic climate change at the Eocene–Oligocene boundary remains a topic of debate. Here, we find a significantly greater sensitivity of Antarctic temperatures to Southern Ocean gateway changes when atmospheric CO2 concentrations are high. In particular, the closure of the Drake Passage (DP) gap is a necessary condition for the existence of ice-free Antarctic conditions at high CO2 concentrations in our coupled climate model. The absence of the Antarctic Circumpolar Current (ACC) is particularly conducive to warm Eocene Antarctic conditions at higher CO2 concentrations, markedly different to previous simulations conducted under present-day CO2 conditions. Antarctic sea surface temperature (SST) and surface air temperature (SAT) warming due to a closed DP gap reach values around ca 5°C and ca 7°C, respectively, for high concentrations of CO2 (above 1250 ppm). In other words, we find a significantly greater sensitivity of Antarctic temperatures to atmospheric CO2 concentration when the DP is closed. The thermal isolation of Antarctica arising from the development of the ACC inhibits a return to the warmer Antarctic and deep ocean conditions resembling the Eocene, even under enhanced atmospheric greenhouse gas concentrations. Figure 1 shows the Antarctic SST warming in response to opening Drake Passage depends on the ambient CO2 level. Many previous studies examining the effect of Drake Passage employ

an ambient radiative balance equivalent to preindustrial atmospheric CO2 concentrations around 280 ppm. Closing DP under these cool conditions leads to Antarctic SST warming (left panel), but the warming anomaly is limited to the north of 60°S and is not circumpolar. Conducting the same experiment under 1250 ppm CO2 leads to significantly enhanced Antarctic warming, and the warming anomaly now reaches to the Antarctic coast and is circumpolar (Fig. 1, right panel). In Figure 2, we have run our model to equilibrium over a period of 7000 years under a range of CO2 values. Several key climate diagnostics are displayed here. Red curves indicate the DP open experiments (DP open), black curves indicate the Drake Passage closed experiments (DP clsd) and dashed curves indicate the difference between the DP closed – DP open cases. In other words, the dashed curve indicates the effect of opening DP alone (for each CO2 value). Although global temperatures are not influenced by DP in our model (Fig. 2, panel a), Antarctic SAT response to closing DP is significantly stronger at higher values of CO2, whereby the average SAT over Antarctica increases by 7°C upon closing Drake Passage at constant high CO2 (Fig. 2, panel b). This is similar for SST and deep ocean temperatures (Fig. 2, panels c and d). Similarly, in Figure 3, Antarctic permanent snow is relatively robust with respect to CO2 when DP is open (red curve, panel a), whereas

Figure 1 Annual Average SST difference DPclsd – DP open under atmospheric CO2 concentration at 280 ppm (left panel) and 1250 ppm (right panel). Warming as a result of closing the DP extends to higher southern latitudes at high CO2 concentrations.

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Figure 2 Dependence of deep ocean and Antarctic climate on atmospheric pCO2 for DP open (red, difference taken with DPopen 280ppm), DPclsd (black, difference taken with DP open 280ppm) and the difference between the two (black, dashed). Globally averaged sea level air temperature (SLAT) change is very similar for DPclsd and DPopen (a), whereas DPclsd exhibits stronger sensitivity to pCO2 than DPopen for annually averaged Antarctic SAT change (b), Antarctic SST change (c) and deep ocean temperature (d). In (b) spatially and annually averaged SLAT has been calculated south of 70°S, whereas in (c) annual averages are taken south of 62° S. Deep ocean temperatures are averaged globally and between 4153m and 5137m depth.

Figure 3 Dependence of factors affecting Antarctic climate on atmospheric pCO2 for DPopen (red) and DPclsd (black). DPclsd exhibits stronger sensitivity to pCO2 than DPopen for permanent snow cover (a), annually averaged sea-ice area (b) and oceanic southward heat transport at 60° S (c). Snow-free conditions are achieved at 1500 ppm pCO2. AABW in DPclsd increases modestly with pCO2 (d). Antarctic permanent snow cover is calculated as the land surface area covered in snow throughout the year.

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permanent snow is significantly more sensitive to CO2 when DP is closed (black curve, panel a), and disappears for CO2 values above 1250ppm.

This is due the reduction of sea ice in a warm ambient climate. The latent heat budget of sea ice acts as a buffer for energy changes south of 60°S in a cool ambient climate.

The Antarctic cryosphere is significantly more sensitive with respect to atmospheric CO2 when Drake Passage is closed and the oceanic thermal isolation of Antarctica arising from the ACC is thus removed. Also, Antarctic temperature changes due to the opening or closing of Drake Passage are significantly greater when ambient atmospheric CO2 is high.

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In conclusion, the effect of Southern Ocean gateways changes is strengthened in a warmer climate, and the interpretation of the results of previous Drake Passage studies where a cool ambient climate is used requires caution.

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CLIMATE AND VEGETATION OF THE SELANDIAN–THANETIAN PALEOCENE SAGWON SECTION, NORTHERN ALASKA (PALEOLATITUDE 85°N) R.A. Spicer1, R.J. Daly2, D. W. Jolley2, A.B. Herman3, A.Ahlberg4 and M. Moiseeva3 1

The Open University, Milton Keynes, MK7 6AA, UK: [email protected]; 2The University of Aberdeen, Aberdeen, AB24 3EU, UK; 3Geological Institute, Russian Academy of Sciences, Moscow 119017, Russia; 4 Lund University, P.O. Box118, SE22100, Lund, Sweden.

The Paleogene of the high paleo-Arctic provides valuable insights into the vegetation dynamics of one of the most climatically sensitive regions on Earth for past, present and future environmental change scenarios. At Sagwon Bluffs, (69°23’N, 148°43’W) Northern Alaska (Fig. 1), combined sedimentological and palynological data (Fig. 2) has identified six depositional sequences within the Prince Creek Formation, Sagwon Member. The base of each sequence is marked by fluvial channel or crevasse-splay sediments, followed by emergent floodplain facies and subsequent deposition of peat mires. These mires gave way to lake sediments with the continued rise in water table and in turn replaced by floodplain water bodies, which in the oldest and youngest sequence contained brackish water algae.


Among the ninety-seven in situ palynomorph taxa recorded are the fungal spore Pesavis tagluensis, and the juglandaceous pollen Caryapollenites imparalis/inelegans indicating an age younger than Danian but older than Upper Paleocene. Taxa characteristic of the PETM are not recorded, suggesting that a Late Paleocene Selandian–Early Thanetian age (ca 61 Ma to 57 Ma; Gradstein et al. 2004) is appropriate. Paleogeographic reconstructions position the area at 85°N at that time (Smith et al. 1981; Ziegler et al. 1983). We have used Correspondence Analysis (Hill 1979) to derive ecological groups representing successional “communities” in the floodplain sediments (e.g. Fig. 3). In some sections this includes a lacustrine group containing Azolla and other freshwater aquatics. These reflect

Locality map for Sagwon section, N. Alaska. Base map from (Mull and Harris 1989).

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Figure 2 Sagwon Section 2 sedimentary log, sequence stratigraphic interpretation and pollen spectra. Section measured at the axis of the Ivishak Anticline (See Fig. 1).

Figure 3 Example of a Correspondence Analysis of Sagwon pollen/spore assemblages interpreted in terms of moisture gradient (Y axis) and seral development (X axis).

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angiosperm understorey and riparian margin taxa (>25 morphotypes) derived in part from earlier Late Cretaceous, more southerly, first occurrences in N.E Russia (e.g. Fig. 4).

Figure 4 Selected leaf forms from Sagwon. a-c) Tiliaephyllum brooksense sp. nov.; d, e) Archaeampelos mulli sp.nov.; f, g) Dicotylophyllum sagwonicum sp.nov.; h) Metasequoia occidentalis. New taxa are described in Moiseeva et al. (in press).

different substrate stabilities and drainage regimes on the Sagwon floodplain. Similar analysis of the seven coal beds shows some originated as lacustrine accumulations of drifted wood and were later subjected to gleysol formation. Autocthonous coals showed a succession from polypodiaceous fern dominance through mid seam Betulaceae, Myricaceae and Fagaceae rich assemblages, to Taxodiaceae-Nyssaceae dominance. This succession is disrupted in some coals by the abundant charcoalified debris indicating that regular wildfires disturbed these mire communities. Leaf, fruit and seed megafossil assemblages characterise more localised communities. These confirm a deciduous taxodiaceous conifer-dominated forest with a rare xeromorphic conifer component (Fokienopsis catenulata and cf. Mesocyparis) that we interpret as evergreens. This mixed coniferous forest supported a diversity of deciduous

Well-developed tree rings (Fig. 5) are, on average, wider than those of the Maastrichtian of northern Alaska (Spicer and Parrish 1990a) (in some cases 5–10mm wide) with only small amounts of latewood (sensu Denne 1989; Mork 1928) indicative of benign summer growth conditions and a pronounced polar light regime.

Overall the composition of the vegetation, coupled with the tree ring characteristics and the abundance and thickness of siliciclastic-free coals suggests high precipitation and a cool temperate thermal regime with mean annual temperatures (MATs) between 3–13°C (Wolfe 1979). However that MAT was likely to have been higher than the 5°C ± 2°C estimated for the Maastrichtian of the region (Spicer and Parrish 1990b) because the Sagwon flora is more diverse and the tree rings generally wider. The presence of some evergreen conifer components suggests winter temperatures were cold enough to depress metabolic rates sufficient to make retention of small xeromorphic leaves an energy-efficient competitive strategy for overwintering. There is no wood anatomical or sedimentological evidence for hard winter freezes. At such high paleolatitudes this is indicative of a weak polar high-pressure system compared to today.

Figure 5 Example of well-preserved Sagwon conifer wood showing distinct growth rings with small amounts of latewood. False rings (arrowed) occur infrequently suggesting summer growing conditions were usually benign throughout the season. Inter-annual variability may indicate some cyclical patterns as evidenced by groups of relatively thin rings and groups of thicker rings. Extended Abstracts

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REFERENCES Moiseeva, M.; Herman, A.B.; Spicer, R.A. 2009 (in press): Late Paleocene flora of Northern Alaska: role of trans-Beringian plant migrations and climate changes in its forming. Paleontological Journal. Denne, M.P. 1989: Definition of latewood according to Mork (1928). IAWA Bulletin 10: 59-62. Gradstein, F.M.; Ogg, J.G.; Smith, A.G. ed. 2005: A Geologic Time Scale 2004. Cambridge, Cambridge University Press. 610 p. Hill, M.O. 1979: Correspondence Analysis – A neglected multivariate method. Applied Statistics 23: 340-354. Mork, E. 1928: Die Qualität des Fichtenholzes unter besonderer Rücksichtnahme auf Schleifund Papierholz. Der Papier-Fabrikant 26: 741747. Mull, C.G.; Harris, E.E. 1989: Road log from Chandalar Shelf (Mile 237.1) to Prudhoe Bay (Mile 414). In C.G. Mull; K.E. Adams ed. Dalton Highway, Yukon River to Prudhoe Bay, Alaska: Bedrock Geology of the eastern Koyukuk basin, central Brooks Range, and east-central Arctic Slope: Alaska: Guide Book

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GB 7. Fairbanks, Alaska, State of Alaska Department of Natural Resources, Division of Geological and Geophysical Surveys. Pp. 101131. Smith, A.C.; Hurley, A.M.; Briden, J.C. 1981: Phanerozoic Palaeocontinental Maps. Cambridge, Cambridge University Press. 102 p. Spicer, R.A.; Parrish, J.T. 1990a: Latest Cretaceous woods of the central North Slope, Alaska. Palaeontology 33: 225-242. Spicer, R.A.; Parrish, J.T. 1990b: Late Cretaceousearly Tertairy palaeoclimates of northern high latitudes: a quantitative view. Journal of Geological Society, London 147: 329-341. Wolfe, J.A. 1979: Temperature Parameters of Humid to Mesic Forests of Eastern Asia and Relation to Forests of Other Regions of the Northern Hemisphere and Australasia. Geological Survey Professional Paper 1106: 137. Ziegler, A.M. Scotese, C.R.; Barrett, S.F. 1983: Mesozoic and Cenozoic paleogeographic maps. In P. Brosche; J. Sündermann ed. Tidal Friction and the Earth's Rotation. Berlin, SpringerVerlag. Pp. 240-252.

Extended Abstracts


BENTHIC FORAMINIFERAL ISOTOPE RECORDS ACROSS THE PETM FROM THE NEW JERSEY COASTAL PLAIN P. Stassen1, E. Thomas2 and R. P. Speijer1 1

Department of Earth Sciences and Enviromental, KULeuven, B-3001, Leuven, Belgium: [email protected]; 2Center for Study of Global Change, Yale University, CT 06520-8109, USA.

High resolution foraminiferal stable isotope records have been published on stratigraphically complete Paleocene–Eocene Thermal Maximum (PETM) deep-sea sequences, but high resolution correlation with neritic records is still in its infancy. Paleoceanographic reconstructions rely on the assumption that foraminiferal stable isotope values reflect ambient environmental conditions, but these reconstructions are complicated by interspecies isotopic offsets that result from distinct microhabitat preferences (test growth in isotopically distinct environments) and vital effects (speciesspecific metabolic variations in isotope fractionation). Ideally, isotope records should be generated using a single foraminiferal species and this is more problematic in shelf sediments than in deep-sea sections, because shallow water environments are inherently more variable, resulting in larger variation in species composition of benthic foraminiferal assemblages. In deep-sea successions, the genus Cibicidoides is traditionally used for isotope analysis. Other commonly used species are Oridorsalis umbonatus, Nuttallides truempyi and Gavelinella beccariiformis and correction factors have been established for these taxa through the analysis of different species within the same sample (Katz et al. 2003). Cores spanning the Paleocene–Eocene boundary, recovered in the New Jersey Coastal Plain (NJ, USA), have been studied at high resolution to reconstruct the environmental effects of the PETM in neritic settings, using stable isotope measurements on benthic foraminifera (e.g. Zachos et al. 2006; John et al. 2008). The New Jersey PETM, as recognized by the occurrence of the characteristic Carbon Isotope Excursion (CIE), corresponds to a clayey interval between the Paleocene glauconitic sands of the Vincentown Formation and the Eocene glauconitic sands and clays of the Manasquan Formation. Sharp lithological contacts occur between Eocene

Extended Abstracts

glauconitic clays and PETM clays, and are interpreted as unconformities on biostratigraphic evidence (e.g. Cramer et al. 1998). We studied benthic foraminiferal assemblages and compared new benthic multispecies foraminiferal isotope data (Wilson Lake; Fig. 1) with published data in order to evaluate the reconstructions of the magnitude of the CIE (necessary for an estimate of the amount of isotopically light carbon released in the oceanatmosphere system). In the Wilson Lake core, Cibicidoides spp. (including C. alleni) only occur in the pre-CIE and post-CIE intervals, while in the CIE interval we observed common specimens of Anomalinoides acutus, which morphologically somewhat resembles C. alleni. Isotopic correction factors between A. acutus and C. alleni are not available in the literature, so we analyzed both taxa in all samples in which they co-occur. Values for the two species are similar, suggesting comparable microhabitat preferences and vital effects. Our values for A. acutus in the CIE interval agree with those for ‘Cibicidoides spp.’ in Zachos et al. (2006). The values for the pre-CIE and post-CIE glauconitic sands in Zachos et al. (2006) agree with our values for C. alleni. We thus assume that Zachos et al. (2006) analyzed Cibicidoides specimens in the pre- and postCIE intervals, and A. acutus specimens (incorrectly identified as Cibicidoides spp.) in the CIE interval. Because the two species do not show interspecies offsets, paleoceanographic interpretations remain valid although the record is based upon two different genera. The Paleocene foraminiferal isotope data are characterized by uniform values and the CIE-onset is marked by an abrupt negative shift of ca 4‰ in carbon isotope values. Benthic foraminiferal isotope values remain constant over the entire clayey CIE interval and show a sudden positive shift of 2‰ across the unconformable transition to the Eocene glauconitic clays. The absence of intermediate

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Figure 1 sections

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Benthic foraminiferal assemblages and stable isotopes from Wilson Lake and Bass River, PETM

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recovery values of the CIE indicates an incomplete record of the upper CIE interval at Wilson Lake. The stable isotope record for Bass River (John et al. 2008; Fig. 1), a slightly deeper site, is probably based on “Cibicidoides” specimens, because the genus occurs in the whole interval (Cramer 1998). Yet, this published Bass River isotope record is still problematic across the base of the CIE. Their “Cibicidoides” values for the CIE interval agree with values for A. acutus and Cibicidoides spp., but the Paleocene values are lighter by about 1‰ in carbon isotope values and are the same as values obtained for G. beccariiformis (Cramer et al. 1999), which is more common and larger than Cibicidoides in that interval. This suggests that John et al. (2008) may have analyzed G. beccariiformis rather than Cibicidoides spp. The oxygen isotope values for these two taxa are similar, but there is no consistent and reliable correction factor known for the carbon isotope values of the two taxa (Katz et al. 2003). The Bass River data suggest at that site the interspecies offset may be about 1‰, thus leading to an underestimate of the amplitude of the CIE by about 1‰. Thus the apparent difference in amplitude of the CIE at Wilson Lake and Bass River may be an artefact of partially incorrect identification of analyzed benthic species. Reassessment of the Bass River and Wilson Lake isotopic records therefore shows that great care and thorough taxonomic understanding of the analyzed benthic foraminifera is needed in order to construct accurate records of the PETM. Finally, the lack of intermediate carbon isotope values, characteristic of the recovery interval, suggests that the deposition of the clayey sediments occurred within the peak interval of the CIE, thus representing a maximum of 70 k.y. on the time scale of Röhl et al. (2007), and possibly less time because it is unknown how

Extended Abstracts

much was eroded during the formation of the unconformity. This compromises any age models and sediment accumulation estimates for this sequence. The suggested accumulation rates for the CIE interval (John et al. 2008) are thus high, but an underestimate. REFERENCES Cramer, B.S. 1998: The late Paleocene thermal maximum, Bass River, NJ: Neritic response to a runaway greenhouse event. MSc thesis, Rutgers University, New Brunswick. Cramer, B.S.; Aubry, M.-P.; Miller, K.G.; Olsson, R.K.; Wright, J.D.; Kent, D.V. 1999: An exceptional chronological, isotopic, and clay mineralogic record of the latest Paleocene thermal maximum, Bass River, NJ, ODP 174AX. Bulletin de la Société Géologique de France 170: 883-897. John, C.M.; Bohaty, S.M.; Zachos, J.C.; Sluijs, A.; Gibbs, S.; Brinkhuis, H.; Bralower, T.J. 2008: North American continental margin records of the Paleocene-Eocene thermal maximum: Implications for global carbon and hydrological cycling. Paleoceanography 23: PA2217. Katz, M.E.; Katz, D.R.; Wright, J.D.; Miller, K.G.; Pak, D.K.; Shackleton, N.J.; Thomas, E. 2003: Early Cenozoic benthic foraminiferal isotopes: Species reliability and interspecies correction factors. Paleoceanography 18: 1024. Röhl, U.; Westerhold, T.; Bralower, T.J.; Zachos, J.C. 2007: On the duration of the PaleoceneEocene thermal maximum (PETM). Geochemistry, Geophysics, Geosystems 8: Q12002. Zachos, J.C.; Schouten, S.; Bohaty, S.; Quattlebaum, T.; Sluijs, A; Brinkhuis, H.; Gibbs, S.J.; Bralower, T.J. 2006: Extreme warming of mid-latitude coastal ocean during the Paleocene-Eocene Thermal Maximum: Inferences from TEX86 and isotope data. Geology 34: 737-740.

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STAR OF THE ANTIPODES: ASTEROCYCLINA IN THE NEW ZEALAND EOCENE C.P. Strong GNS Science, 1 Fairway Drive, Lower Hutt, New Zealand: [email protected]

INTRODUCTION The starlike image of Asterocyclina speighti in the CBEP 2009 logo aptly symbolises the conference theme of ancient greenhouse climates and the warm Paleogene in New Zealand (NZ). This foraminifer, with its distinctive shape and large size, is tangible and striking evidence that NZ’s Eocene sea surface temperatures were significantly warmer than today’s yearly average of ca 13 to 18ºC. Prior to advent of sophisticated and sensitive geochemical and isotopic techniques for determining paleotemperature, A. speighti, together with other elements of the NZ fossil biota, provided the main indicators of past climates. Early paleontological work in NZ tended to focus on the age correlation utility of fossils within individual specialists’ groups (e.g., molluscs), although paleoenvironmental implications were occasionally acknowledged. Adams et al. (1990) and Hornibrook (1992) produced probably the first major syntheses of NZ Cenozoic climate history employing a range of paleontological criteria. They cited

Figure 1 Asterocyclina speighti, showing complex internal structure visible in axial and longitudinal thin sections. Note the large medial primary chambers, and several layers of smaller secondary chambers, stellate rays and pillars (From Cole 1962).

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Eocene occurrences of Asterocyclina and Amphistegina (visually uninspiring!), along with Lingula, various warm-water molluscs, scleractinian corals, bryozoans and widespread Nypa (mangrove palm), Cocos (coconut), and other semitropical/tropical plants, as evidence that much of the NZ Eocene was warmer than present day, despite the country then being at a paleolatitude of ca 55ºS. Adams et al. (1990) made no specific interpretation of Eocene paleotemperature, but suggested a range of 2028oC for sea surface temperatures during early Eocene to middle Miocene. Hornibrook (1992) interpreted sea surface temperatures of up to ca 23ºC for the lower and Middle Eocene, brief cooling in the late Middle Eocene, and warming, to ca 21ºC, for the remainder of the epoch. Recent geochemical results from the mid-Waipara River section, North Canterbury, indicate Eocene paleotemperatures of >30ºC (Hollis et al. 2009), considerably higher than any of the earlier estimates. LARGER FORAMINIFERA Asterocyclina, and also Amphistegina, along with a score or more other genera, many only distantly related, comprise a loosely defined group of “larger” foraminifera, whose main characteristics are: •

Adult diameter 2 mm or more,

•

Simple to modified, compressed lenticular to discoidal shape, often with external ornament,

•

Complex internal structure, with primary and secondary chambers, canal systems and pillars, often necessitating thin sections for taxonomic study (Fig. 1),

•

Contain symbiotic algae, restricting them to shallow (25ºC mean summer isotherm, and cannot reproduce in waters colder than 17-20ºC.

Extended Abstracts


THE NZ ASTEROCYCLINA SPECIES Two species of Asterocyclina are recognised in NZ Eocene strata (Cole 1962, 1967): Asterocyclina speighti (Chapman) External description: Test large (to 4 mm), robust, usually stellate or stellate-ridged, with four to six (commonly five or six) rays; moderately raised central area and coarsely pustular ornament, denser on ribs than in the interrib areas (Figs 1 & 2). Stouter and more coarsely ornamented than A. hornibrooki. NZ Stratigraphic Range: Waipawan to Mangaorapan (or Heretaungan?) Stages, 55.5– 49.5 (or 46.2?) Ma. Topotypes from Whites Creek dated as Mangaorapan by co-occurrence with Morozovella subbotinae, Globanomalina wilcoxensis, Subbotina cf. velascoensis, and Vaginulinopsis marshalli. Asterocyclina hornibrooki (Cole) External description: Test large (to 5 mm), often stellate or stellate-ridged, with four to seven rays; central area slightly raised with coarse papillae; finer, densely papillate ornament evenly covers the remainder of the test (Fig. 3). Is somewhat more delicate, flattened and finely ornamented than A. speighti. NZ Stratigraphic Range: Runangan Stage (Late Eocene), 36.0–34.3 Ma. Dated by cooccurrence with Globigerapsis index, Bolivina pontis, and Cibicides verrucosus.

Figure 2 Asterocyclina speighti, topotypes showing a range of shape variation. Note that specimens may show radiating ridges but lack stellate branches on periphery. Extended Abstracts

HISTORY OF NZ ASTEROCYCLINA STUDIES Asterocyclina speighti was the first species discovered. R. Speight (1928), a noted Canterbury geologist, recorded the species in beds along the Eyre River in mid Canterbury, listing it, without formal description, as Orthophragmina sp. In 1932, F. Chapman, an Australian micropaleontologist, formally described Speight’s material from thin sections as Discocyclina speighti, and also identified six other species of nummulitids and orbitoids - all later considered to be A. speighti - in the same sample. H.J. Finlay (1946) published a strongly critical review of Chapman’s taxonomy of numerous south Canterbury foraminifera, and, amongst other revisions, placed speighti within Asterocyclina. Finlay observed on Chapman’s identifications: “These slides and identifications are in such a state that in most cases it would not be worth the time and trouble to make out a corrected list of the identifications….; it is enough to state that over 90% are specifically erroneous, and far more than half, generically so”. And, about A. speighti, “Chapman’s material was so badly decorticated and glauconitised that all surface features were obscured and almost all specimens badly broken; his microphotographs show no significant internal differences.” The discoverer of A. hornibrooki is apparently unrecorded, but Hornibrook (1958), in a discussion of NZ foraminiferal zones, called attention to the Late Eocene occurrence of the younger Asterocyclina species, assigning it to undifferentiated Discocyclinidae. Later, he submitted a suite of NZ specimens to W. Storrs Cole, then the world authority on larger foraminifers, for taxonomic analysis. Type specimens of A. hornibrooki were described (Cole 1967) from the Totara Limestone (Runangan; late Eocene), at Fortification Hill, near Oamaru, in North Otago. The species was recognised as differing in several respects from A. speighti, and initially, was tentatively assigned (Cole 1962) to the previously named tropical species, A. matanzensis and A. praecipua. W. Storrs Cole (1962, 1967), produced the definitive studies of New Zealand Asterocyclinas. He confirmed Finlay’s diagnosis of A. speighti and defined the new

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species, A. hornibrooki, for the specimens from the upper Eocene. Cole’s reports represent the last significant work on the taxonomy of NZ Asterocyclina. NEW ZEALAND ASTEROCYCLINA OCCURRENCES Geographically, distribution of Asterocylina in New Zealand is apparently limited to a few outcrops in Canterbury, North Otago and Chatham Island, and consequently the genus has played little role in biostratigraphic correlation. A. speighti has been recorded in tuffaceous sandstones in the Eyre RiverWhites Creek area (most earlier exposures seem to have disappeared), from conglomeratic sandstones near Waihao (Riddolls 1966), some 200 km to the southwest, and from the Matanginiu Limestone on the Chatham Islands (Hornibrook 1992). Records of A. hornibrooki are from volcanic-associated limestones within the vicinity of its type locality at Fortification Hill, near Oamaru, North Otago (Lee et al. 1997). Hornibrook et al. (1989) noted that, for large foraminifers in general, “Their occurrence in New Zealand in the Cenozoic is controlled both by past climate and by the rather sporadic development and preservation of suitable facies. It is uncertain whether their discontinuous presence is due to fortuitous preservation or to a series of immigrations and extinctions.” The widespread distribution of offshore facies in the NZ Eocene bearing a persistent warm-climate geochemical signature (Hollis et al. 2009), suggests the former

explanation is the more likely, and the common occurrence in volcanic sedimentary associations suggests that relatively short-lived volcanic edifices often provided the necessary shallow water habitat. ACKNOWLEDGMENTS Many thanks to Daphne Lee for reviewing the draft abstract and for providing the sample with the well preserved Asterocyclina hornibrooki specimens figured here, to Hugh Morgans for his comments on the draft manuscript, and to Andy Gray for producing the figures. REFERENCES Adams, C.G.; Lee, D.E.; Rosen, B. 1990: Conflicting isotopic and biotic evidence for tropical sea-surface temperatures during the Tertiary. Palaeogeography, Palaeoclimatology, Palaeoecology 77: 289313. Chapman, F. 1932: On a rock containing Discocyclina and Assilina, found near Mt. Oxford, South Island, New Zealand. Records of the Canterbury Museum (NZ) 3: 483-489. Cole, W.; Storrs 1962: Asterocyclina from New Zealand and the Chatham Rise. Bulletin of American Paleontology 44: 343-357. Cole, W.; Storrs 1967: Additional data on New Zealand Asterocyclina (Foraminifera). Bulletin of American Paleontology 52: 5-18. Finlay, H.J. 1946: Microfaunas of the Oxford Chalk and Eyre River Beds. Transactions of the Royal Society of New Zealand 76: 237-243. Hollis, C.J.; Handley, L.; Crouch, E.M.; Morgans, H.E.G.; Baker, J.A.; Creech, J.; Collins, K.S.; Gibbs, S.J.; Huber, M.; Schouten, S.; Zachos, J.C.; Pancost, R.D. 2009: Tropical sea temperatures in the high-latitude South Pacific during the Eocene. Geology 37: 99-102. Hornibrook, N. deB. 1958: New Zealand Upper Cretaceous and Tertiary foraminiferal zones and some overseas correlations. Micropaleontology 4: 25-38.

Figure 3 Asterocyclina hornibrooki, from type area, showing a range of shape variation. Note that specimens may show radiating ridges but lack stellate branches on periphery.

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Hornibrook, N. deB. 1992: New Zealand Cenozoic marine paleoclimates: a review based on the distribution of some shallow water and terrestrial biota. In Tsuchi, R.; Ingle, J.C. ed. Pacific Neogene, environment, evolution and events. University of Tokyo Press. Pp. 83-106. Hornibrook, N. deB.; Brazier, R.C.; Strong, C.P. 1989: Manual of New Zealand Permian to

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Pleistocene foraminiferal biostratigraphy. New Zealand Geological Survey Paleontological Bulletin 56: 175 p. Lee,

D.E.; Scholz, J.; Gordon, D. 1997: Paleoecology of a late Eocene mobile rockground biota from North Otago, New Zealand. Palaios 12: 568-581.

Extended Abstracts

Riddolls, B.W. 1966: Note on the occurrence of Asterocyclina speighti (Chapman). New Zealand Journal of Geology and Geophysics 5: 471-473. Speight, R. 1928: The geology of Vine Hill and neighbourhood. Transactions of the Royal Society of New Zealand 58: 104-131.

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ORGANIC GEOCHEMISTRY OF PALEOCENE SEDIMENTS FROM THE MIDWAIPARA RIVER SECTION, NEW ZEALAND K. W. R. Taylor1, L. Handley1, C.J. Hollis2, H.E.G. Morgans2, E. M Crouch2, S. Schouten3 and R. D. Pancost1 1

Organic Geochemistry Unit, School of Chemistry, University of Bristol, Bristol, UK: [email protected]; GNS Science, P.O. Box 30-368, Lower Hutt, New Zealand; 3Royal Netherlands Institute for Sea Research, P.O. Box 59, 1790 AB Den Burg, Texel, The Netherlands.

2

The Paleocene is an intriguing interval for paleoclimate investigations, being bounded by the global biological catastrophe that occurred at the Cretaceous–Paleogene (K–Pg) boundary and the dramatic warming associated with Paleocene–Eocene Thermal Maximum, which led into the Early Eocene Climatic Optimum. An intriguing aspect of the Paleocene is a pronounced positive excursion in δ13C observed in marine carbonate at ca 59–56 Ma (Kurtz et al. 2003). This Paleocene Carbon Isotope Maximum (PCIM) indicates accelerated rates of carbon burial and appears to be associated with oxygen isotopic evidence for cooling (Zachos et al. 2001). Correlative increases in organic content and siliceous microfossil abundance offshore New Zealand suggest an associated increase in oceanic productivity (Hollis 2002). It is uncertain if this productivity signal is a local response to global cooling or is due to other factors, such as regional oceanographic (e.g. upwelling) changes. We have employed the use of geochemical proxies and biomarkers to interrogate Paleocene sedimentary records at the basinal scale in New Zealand and also at Walvis Ridge with the aim of reconstructing climatic conditions at this period in geological history. Here, we focus on the Waipara Greensand Formation in the mid-Waipara River section, Canterbury Basin, New Zealand. The section has long been recognised as an exceptional geological record for paleoclimate studies as it contains a near complete succession of neritic to upper bathyal sediments from the Late Cretaceous to Middle Eocene. The 70 m thick Waipara Greensand appears to span the entire Late Paleocene. The base is marked by the first occurrence of calcareous nannofossil index species Chiasmolithus bidens (upper NP4, 60.7 Ma; Berggren et al. 1995) and the top contains the first occurrence of the dinoflagellate genus Apectodinium, an event that occurs very close

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to the Paleocene–Eocene boundary in New Zealand (Cooper 2004). The lower 20 m of the formation is correlated with the early Late Paleocene (~60 Ma) based on the concurrence of dinoflagellate cyst species Alterbidinium pentaradiatum and Deflandrea foveolata (Cooper 2004). The upper 10 m of the formation is characterised by high total organic carbon (TOC >0.5) and positive organic δ13C (>-25 per mil). These features suggest a correlation with the Late Paleocene Waipawa Formation — an organic-rich mudstone that is found in most of New Zealand’s sedimentary basins, which is invariably enriched in δ13C (Hollis 2006). The unit appears to have been deposited during the PCIM, although initial studies indicate that the PCIM is a longer duration global phenomenon that incorporates localised deposition of organic-rich sediments during an early phase (Hollis et al. 2005). Our study examines variation in biomarkers and organic geochemical climate proxies through the upper part of the Waipara Greensand with an emphasis on elucidating the climatic conditions associated with deposition of this organic-rich interval. Thirty sediment samples were collected at 0.5– 2 m intervals in the upper 20 m of the Waipara Greensand and at 2–9 m intervals in the lower part. Sample splits of ~20 g were used for geochemical analysis. Total lipid extracts were obtained by Soxhlet solvent extraction of freeze-dried and ground samples. The total lipid extracts were separated into neutral apolar, neutral polar and acid fractions (Fig. 1) by a series of solid phase extractions and flash chromatography techniques. These fractions were then analysed by GC-MS. The analyses yielded a suite of algal, higher plant and bacterial organic biomarkers that allow reconstruction of Paleocene climate and ecological conditions.

Extended Abstracts


Figure 1 Typical partial gas chromatogram of (A) apolar (m/z 71) and (B) acid (total ion current) lipid fractions from Mid-Waipara.

The biomarkers and lipids present in the sediments allow further interrogation of the sedimentary record. Parameters related to higher plant leaf wax n-alkane and n-alkanoic acid distributions all indicate a similar pattern, particularly in the upper 10 m of the section, characterised by high TOC and enriched bulk organic δ13C. The n-alkane carbon preference index (CPI) increases from 0.56 at ca 80 m to 1.58 at 128 m, whilst the odd-over-even predominance (OEP) similarly increases from

0.44 to 1.96, both indicating a greater input of terrigenous material to the latest Paleocene sediments. The average chain length (C24-C32) also shifts towards higher chain preference at around 120–130 m, possibly indicating a shift in the plant communities providing the source of terrestrial material to the sediment. These trends are mirrored in the acid fraction by the n-alkanoic acids — again the CPI increases from 1.87 at ca 80 m to 2.87 at 128 m, the even-over-odd (EOP) predominance shifts

Figure 2 Trends in geochemical parameters in the Waipara Greensand (column 2), mid-Waipara River, Canterbury, New Zealand: (A) Total organic carbon, (B) δ13C (total organic), (C) High Molecular Weight (C26-C32) n-alkane abundance, (D) Stigmastanol abundance. Extended Abstracts

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from 1.79 to 2.89, and the ACL again shows a general shift towards higher chain fatty acids. The evidence for increased terrestrial material around 120–130 m from n-alkyl compound ratios is corroborated by an increase in stigmastanol and high molecular weight (C26C32) alkane abundances throughout the same interval (Fig. 2), and is also consistent with terrestrial palynomorph assemblages. Enhanced transport of allochthonous material to the sediment at this interval may indicate a sea level fall bringing the site closer to shore, or increased terrestrial run-off into the marine system. The sediments of the upper 10 m of the section are also characterised by elevated abundances of archaeol and biphytane diols, potentially of methanogen origin, and a measurable abundance of GDGTs. In total, these records confirm that dramatic environmental changes occurred in the latter part of the Paleocene. Ongoing work reveals similar changes in other Southern Ocean settings and at Walvis Ridge; further comparisons will be used to examine the relationships between cooling, carbon cycle perturbations and oceanographic conditions. REFERENCES Berggren, W.A.; Kent, D.V.; Swisher, C.C.; Aubry, M.-P. 1995: A revised Cenozoic geochronology and chronostratigraphy. In Berggren, W.A.; Kent, D.V.; Aubry, M.-P.; Hardenbol, J. ed. Geochronology, time scales and global stratigraphic correlation. SEPM

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(Society for Sedimentary Geology), Special Publication 54. Pp. 129-212. Cooper, R.A. ed. 2004: New Zealand geological timescale. GNS Monograph 22. Institute of Geological and Nuclear Sciences, Lower Hutt. 284 p. Hollis, C.J. 2002: Biostratigraphy and paleoceanographic significance of Paleocene radiolarians from offshore eastern New Zealand. Marine Micropaleontology 46: 265316. Hollis, C.J.; Dickens, G.R.; Field, B.D.; Jones, C.J; Strong, C.P. 2005: The Paleocene–Eocene transition at Mead Stream, New Zealand: a southern Pacific record of early Cenozoic global change. Palaeogeography, Palaeoclimatology, Palaeoecology 215: 313343. Hollis, C.J.; Field, B.D.; Crouch, E.M.; Sykes, R. 2006: How good a source rock is the Waipawa (black shale) Formation beyond the East Coast Basin?: an outcrop-based case study from Northland, New Zealand. New Zealand Petroleum Conference (2006: Auckland, NZ). Pp. 8. Kurtz, A.C.; Kump, L.R.; Arthur, M.A.; Zachos, J.C.; Paytan, A. 2003: Early Cenozoic decoupling of the global carbon and sulfur cycles. Paleoceanography 18: 1090, doi:10.1029/2003PA000908. Zachos, J.C.; Pagani, M.; Sloan, L.C.; Thomas, E.; Billups, K. 2001: Trends, rhythms, and aberrations in global climate 65 Ma to present. Science 292: 686-693.

Extended Abstracts


THE CRETACEOUS–PALEOGENE (K–PG) CARBON CYCLE E. Thomas1,2, A, Ridgwell3, L. Alegret4 and D. N. Schmidt5 1

Center for Study of Global Change, Department of Geology and Geophysics, Yale University, CT 06520-8109 New Haven, USA: [email protected]; 2Department of Earth and Environmental Sciences, Wesleyan University. CT 06459-0139 Middletown, USA; 3Bristol Initiative for the Dynamic Global Environment, School of Geographical Sciences, University of Bristol, UK; 4Departamento de Ciencias de la Tierra, Universidad de Zaragoza. 50009 Zaragoza, Spain; 5Department of Earth Sciences, University of Bristol, UK.

INTRODUCTION During the Cretaceous–Paleogene (K–Pg) mass extinction, carbonate-secreting oceanic surface dwellers (calcareous nannoplankton, planktic foraminifers) suffered extreme extinction (e.g. Kiessling and Claeys 2001; d’Hondt 2005), whereas deep oceanic bottom dwellers (benthic foraminifera) were not significantly affected (Culver 2003; Thomas 2007). At the same time as the extinction of planktic foraminifera and calcareous nannoplankton, the vertical (planktic-benthic, i.e. surface-bottom) gradient in carbonate carbon isotope values (or ∆δ13C) collapsed rapidly, and remained so for several hundred thousands of years (e.g. Hsü and MacKenzie 1985; Zachos and Arthur 1986). This collapse has been observed in cores recovered from the Pacific, Atlantic, Indian, Southern and Tethys Oceans, and in land sections bordering these oceans (e.g. Kiessling and Claeys 2001). The global collapse of the carbon isotope gradient was first explained as caused by a prolonged (several hundred k.y.) nearly complete cessation of primary productivity in the oceans (‘Strangelove Ocean’; Hsü and MacKenzie 1985). An alternate explanation proposed that the collapse of the carbon isotope gradient was caused by the collapse of the biological pump, i.e. the transport of organic matter from surface to bottom, possibly due to the extinction of pellet producing zooplankton (‘Living Ocean Model’; d’Hondt et al. 1998). According to this ‘Living Ocean Model’, the primary productivity in the oceans may have recovered within ca 10 k.y., similar to the recovery time proposed for the terrestrial biosphere (Beerling et al. 2001). The post-extinction primary producers were dominantly non-calcifying taxa according to the ‘Living Ocean Model’, and the recovery of diversity, especially by calcifiers, took much longer, up to several million years (d’Hondt et al. 1998; Coxall et al. 2006). Extended Abstracts

The two models (Fig.1) thus propose very different causes for the collapse of the vertical isotope gradient, with the Strangelove Ocean Model arguing for a lack of primary productivity of organic matter, and the Living Ocean Model for a lack of transport of organic matter from surface to sea floor, but both models invoke an almost complete cessation of the supply of food to the benthic community. In the present oceans, biota (including foraminifera) in the extremely food-starved deep-sea environment are very strongly influenced by primary productivity in the overlying surface waters (bentho-pelagic coupling, e.g. Gooday 2003; Jorissen et al. 2007). Therefore it seems extremely unlikely that Maastrichtian benthic foraminiferal assemblages would have been able to survive such a prolonged period without food, showing no significant net extinction and only transient changes in assemblage composition. The transient changes in diversity and abundance of taxa vary geographically, indicating a decrease in food supply at some locations (e.g. the Southern Ocean; Thomas 1990), an increase in food delivery at other locations (e.g. in the Pacific Ocean; Alegret and Thomas 2005, this volume). In addition, neither a collapse of primary productivity nor of the biological pump can explain the reversed vertical carbon isotope gradient (rather than just a collapsed gradient) after the extinction, as observed in the southern Atlantic and western Tethys Oceans (e.g. Hsü and MacKenzie 1990; Kaiho and Lamolda, 1999; Thomas et al. 2007). CARBON ISOTOPE RECORDS ACROSS A MASS EXTINCTION The carbon isotopic composition of marine calcifiers is assumed to reflect the isotopic composition of total dissolved inorganic carbon (DIC) in surface (bulk, planktic foraminifera) and deep (benthic foraminifera) oceans (e.g. Maslin and Swann 2006), and this 145


Figure 1 Maps of modelled distribution of δ13C of Dissolved Inorganic Carbon (DIC) in the surface waters of the oceans, in modern (upper) and Maastrichtian (lower) times, superimposed on figure showing the vertical distribution of δ13C in DIC as a mean value for the Pacific Ocean. Note that modern values (upper axis, blue) are plotted on a different scale than the Maastrichtian values (lower axis, green). The blue line shows the modern vertical profile, the interrupted green line shows the Maastrichtian profile at background Maastrichtian primary productivity, the thick green line shows the profile for a reduction of the biological pump by 40%, and the dotted green line shows the profile (including a ‘reversed gradient’) for a complete collapse of the biological pump (‘Strangelove Ocean’).

assumption is at the base of both the ‘Strangelove Ocean’ and ‘Living Ocean’ model. Most taxa, however, including calcareous nannoplankton, planktic foraminifera and benthic foraminifera, precipitate carbonate out of isotopic equilibrium with sea water (the so-called ‘vital effect’). Hence, one or few cross-calibrated species are usually analyzed for their respective isotopic composition (e.g. Katz et al. 2003). Across the K–Pg extinction, carbon isotopic values in a single species of benthic foraminifera can readily be analyzed, but this is impossible for the planktic taxa because all planktic calcifiers suffered mass extinction. In the Cretaceous, bulk carbonate was dominated by calcareous nannoplankton, but bulk carbonate deposited after the extinction (in the Danian) consists of diverse components, including calcareous dinocysts which have a very light carbon isotope signature (e.g. Gardin and Monechi 1998; Minoletti et al. 2005).

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Post-extinction planktic foraminifer assemblages were dominated by biserial and triserial forms (e.g. Koutsoukos 1996; Olsson et al. 1999), and Recent as well as Cenozoic biserial planktic taxa are characterized by a light carbon isotopic signature (e.g. Sexton et al. 2006; Smart and Thomas 2006). In addition, diagenesis in low-carbonate sediments commonly affects isotope records. The interpretation of the planktic carbon isotope record thus may be complex, and this record does not necessarily reflect values of dissolved inorganic carbon in surface ocean waters. Therefore the collapse of the vertical carbon isotopic gradient does not necessarily accurately reflect an almost complete collapse of marine primary productivity or the biological pump. DR. STRANGELOVE MEETS GENIE We argue that the apparent paradox between the collapse in the carbon isotope gradient and Extended Abstracts


the lack of extinction of deep-sea benthic foraminifera is not solely due to such vital effects. Traditional geochemical box models (e.g. Kump 1991) have been applied to model the carbon isotope gradient assuming the collapse in gradient was explained by the operation of the biological pump only. Such models, however, cannot incorporate the role of the ‘solubility pump’ in the oceans. The ‘solubility pump’ imprints an ‘inverted’ isotopic gradient, opposite in sign to that generated by the biological pump (i.e. it causes the benthic signal to be more positive than the planktic one). We assessed the relative importance of solubility vs. biological pump controls on observed ∆δ13C changes across the K–Pg boundary using an Earth system model (GENIE-1, Ridgwell et al. 2007; Ridgwell and Hargreaves 2007), simulating the impact of a range of different strengths of the biological pump on vertical carbon isotope gradients (Fig.1). Late Cretaceous ocean circulation and stratification, latitudinal sea-surface temperature gradients, and productivity distributions all weight the solubility and biological pump δ13C controls differently compared to the modern ocean, and therefore we employed a Maastrichtian continental and geochemical configuration of the GENIE-1 model. We found that a collapsed surface-todeep δ13C gradient is achieved without a complete shutdown of the biological pump. At a global average, a biological pump no weaker than ca 30% of that estimated for the Late Cretaceous is consistent with ∆δ13C ≈ 0.0‰. For the Pacific basin, export production of ca 40% could persist (Fig. 1). We conclude that a Strangelove ocean did not occur in the aftermath of the bolide impact at the K–Pg boundary, and that the flux of organic matter reaching the deep ocean persisted at a rate no less than ca 30–40% (averaged globally) of the pre-impact value. The persistence of organic matter transport to depth is consistent with the observed lack of significant benthic extinction. ACKNOWLEDGEMENTS ET acknowledges funding by NSF Grant OCE 720049. LA acknowledges support from a “Ramon y Cajal” research grant from the Spanish Ministry of Science and Technology and the European Social Fund.

Extended Abstracts

REFERENCES Alegret, L.; Thomas, E. 2005: Cretaceous/Paleogene boundary bathayl paleoenvironments in the central North Pacific (DSDP Site 465), the Northwestern Atlantic (ODP Site 1049), the Gulf of Mexico and the Tethys: The benthic foraminiferal record. Palaeogeography, Palaeoclimatology, Palaeoecology 224: 53-82. Alegret, L.; Thomas, E. 2009: The Pacific Ocean: no Strangelove Ocean after the CretaceousPaleogene impact. GNS Science Miscellaneous Series 18: 10-13. Beerling, D. J.; Lomax, B. H.; Upchurch, G. R.; Nichols, D. J.; Pillmore, C. J.; Handley, L. L.; Scrimgeous, C. M. 2001: Evidence for the recovery of terrestrial ecosystems ahead of marine primary production following a biotic crisis at the Cretaceous-Tertiary boundary. Journal of the Geological Society, London 158: 737-740. Coxall, H. K.; d’Hondt, S.; Zachos, J. C., 2006: Pelagic evolution and environmental recovery after the Cretaceous-Paleogene mass extinction. Geology 34: 297-300. Culver, S. J. 2003: Benthic foraminifera across the Cretaceous-Tertiary (K-T) boundary: a review. Marine Micropaleontology 47: 177-226. D’Hondt, S., 2005: Consequences of the Cretaceous/Paleogene mass extinction for marine ecosystems. Annual Reviews of Ecology and Systematics 36: 295-317. D’Hondt, S.; Donaghay, P.; Zachos, J.C.; Luttenberg, D.; Lindinger, M. 1998: Organic carbon fluxes and ecological recovery from the Cretaceous-Tertiary mass extinction. Science 282: 276-279. Gardin, S.; Monechi, S., 1998: Palaeoecological change in middle to low latitude calcareous nannoplankton at the Cretaceous/Tertiary boundary. Bulletin de la Societé géologique de France 169: 709-723. Gooday, A.J. 2003: Benthic foraminifera (Protista) as tools in deep-water palaeoceanography: environmental influences on faunal characteristics. Advances in Marine Biology 46: 1-90. Hsü, K.J.; McKenzie, J. 1985: A “Strangelove Ocean” in the earliest Tertiary. Geophysical Monographs 32: 487-492. Hsü, K. J.; McKenzie, J. 1990: Carbon-isotope anomalies at era boundaries: global catastrophes and their ultimate cause. Geological Society of America Special Paper 247: 61-70.

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Jorissen, F.J.; Fontanier, C.; Thomas, E. 2007: Paleoceanographical proxies based on deep-sea benthic foraminiferal assemblage characteristics. In Hillaire-Marcel, C.; de Vernal, A. ed. Proxies in Late Cenozoic Paleoceanography: Pt. 2: Biological tracers and biomarkers. Elsevier. Pp. 263-326. Kaiho, K.; Lamolda, M. A. 1999: Catastrophic extinction of planktonic foraminifera at the Cretaceous/Tertiary boundary evidenced by carbon and oxygen isotopes at Caravaca, Spain. Geology 37: 355-358. Katz, M. E.; Wright, J. D.; Katz, D. R.; Miller, K. G.; Pak, D. K.; Shackleton, N. J.; Thomas, E. 2003: Early Cenozoic benthic foraminiferal isotopes: species reliability and interspecies correction factors. Paleoceanography 18: 1024, doi: 10.1029/2002 PA000798 Kiessling, W.; Claeys, P. 2001: A geographic database to the KT boundary. In Buffetaut, E., and Koeberl, C., ed. Geological and biological effects of impact events: Berlin, SpringerVerlag. Pp. 33-140. Koutsoukos, E.A.M. 1996: Phenotypic experiments into new pelagic niches in early Danian planktonic foraminifera: after-math of the K/T boundary event. In Hart, M.B. ed. Biotic Recovery from Mass Extinction Events. Geological Society Special Publication 102. Pp. 319-335. Kump, L. R. 1991: Interpreting carbon-isotope excursions: Strangelove Oceans. Geology 19: 299-302. Maslin, M. A.; Swann, G. E. A. 2006: Stable isotopes in marine sediments. In Leng, M. J., ed. Isotopes in Palaeoenvironmental Research, Springer (Dordrecht, Netherlands), Ch. 6. Pp. 227-290. Minoletti, F.; de Rafaelis, M.; Renard, M.; Gardin, S.; Young, J. 2005: Changes in the pelagic fine fraction carbonate sedimentation during the Cretaceous-Paleocene transition: contribution of the separation technique to the study of the Bidart section. Palaeogeography,

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Palaeoclimatology, Palaeoecology 216: 119137. Olsson R. K.; Hemleben C.; Berggren W. A.; Huber, B. T. 1999: Atlas of Paleocene Planktonic Foraminifera. Smithsonian Contributions to Paleobiology 85: 252 p. Ridgwell, A.; Hargreaves, J. 2007: Regulation of atmospheric CO2 by deep-sea sediments in an Earth System Model, Global Biogeochemical Cycles 21, doi:10.1029/2006GB002764. Ridgwell, A. J.; Hargreaves, J. C.; Edwards, N. R.; Anna, J. D.; Lenton, T. M.; Yool, A.; Marsh, R.; Watson, A. J. 2007: Marine geochemical data assimilation in an efficient Earth system model of global biogeochemical cycling. Biogeosciences 4: 87-104. Sexton P. F.; Wilson P. A.; Pearson, P. N. 2006: Palaeoecology of late middle Eocene planktic foraminifera and ecological implications. Marine Micropaleontology 60: 1-16. Smart, C. W.; Thomas, E. 2006. The enigma of early Miocene biserial planktic foraminifera, Geology 34: 1041-1044. Thomas, E., 1990: Late Cretaceous-early Eocene mass extinctions in the deep sea. Geological Society of America Special Publication 247: 481- 495. Thomas, E. 2007: Cenozoic mass extinctions in the deep sea; what disturbs the largest habitat on Earth? In Monechi, S.; Coccioni, R.; Rampino, M. ed. Large Ecosystem Perturbations: Causes and Consequences. Geological Society of America Special Paper 424. Pp. 1-24. Thomas, E.; Schmidt, D.; Kroon, D.; Alegret, L.; Bernaola, G.; Lohmann, K. C.; Monechi. S.; Roehl, U.; Westerhold. T. 2007; Effects of the end Cretaceous asteroid impact, Walvis Ridge, SE Atlantic Ocean. ICP9, China, September 37, 2007: 96. Zachos, J. C.; Arthur, M. A., 1986: Paleoceanography of the Cretaceous-Tertiary boundary event: inferences from stable isotope and other data: Paleoceanography 1: 5-26.

Extended Abstracts


PALEOGENE NANNOPLANKTON AND DINOFLAGELLATE CYST BIOSTRATIGRAPHY OF THE NORHERN PRICASPIAN DEPRESSION O.N. Vasilyeva1 and V.A. Musatov2 1

Institute of Geology and Geochemistry, Ural Branch of RAS: [email protected]; 2Lower Volga Institute of Geology and Geophysics: [email protected]

INTRODUCTION One of the most complete northern marginal Peritethys epicontinental sections of the Paleogene was cored in the Novouzensk key borehole, near the northern boundary of the Pricaspian Depression. The borehole penetrated a 526 m sequence of marine calcareous-terrigeneous and siliceousterrigeneus sediments that span the Danian– Lutetian interval. The Novouzensk key borehole is located in the right bank of Bolshoy Uzen River, near Novouzensk t. of the Saratov region, Russia (Fig. 1). A study of nannoplankton and dinocyst assemblages has been completed and the results presented below (Fig.2). LITHOSTRATIGRAPHY Paleogene sediments span the white chalk rocks of Maastrichtian. The units are, in ascending order:


Paleocene (Danian), Algay Formation 926–917 m: greenish-grey, dense marl. 917–892 m: in the lower part, greenish-grey, sandy marl; in the upper part, greenish, whitish grey, dense, slightly calcareous clay. Danian, Cygan Formation 892–856 m: black, sandy glauconite clay with calcareous sandstone layers. Selandian, Syzran Formation and Lower Syzran Subformation 856–789 m: dark grey sandstone at the base; upper part, dark grey and black silicified clay with slightly sandy clay and sandstone layers. Upper Syzran Subformation 789–721 m: grey sandstone at the base; upper, alternating dark grey sandy clay and sandstone layers. Clay is slightly calcareous.

The location of the Novouzensk key borehole.

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Paleocene–Eocene (Ypresian), Novouzensk Formation

N. perfectus, Ch. danicus, C. tenuis and rare specimens of B. sparsus.

721–523 m: in the lower part, dark grey, black, sandy clay with thin sand and silt lenses. Clay includes plant remains. Clay is slightly calcareous in the lower part (25 m) of this unit. The upper part consists of dark grey, sandy, micaceous, non-calcareous clay with thingrained sandstone layers.

The silicious clay of the Lower Syzran Subformation (856–789 m) contains abundant nannofossils representing the Zone Neochiastozygus junctus (upper zone NP4 Coccolithus robustus or CP3 Ellipsolithus macellus). Abundant specimens of N. junctus, and a single species C. consuetus, are present. Zonal species E. macellus, typical for West Europe and oceanic zonal schemes, and also E. distichus occur only rarely.

Eocene (Ypresian), Bostandyk Formation 523–486 m: alternating dark grey silt clay and clay glauconite-quartz sandstone layers with molluscan remains. There are slightly calcareous rocks in the upper part. 486–405 m: in the lower part, alternating dark green, sandy, calcareous clay, fine-grained glauconite sandstone, and marl beds; in the upper part, alternating silt and sandy-clay rocks. Clay includes molluscs, sponge spicules, fish scales. This interval is irregularly calcareous. Lutetian, Kopterek Formation 405–400 m: glauconite clay calcareous sandstone at the base, grading into greenish, whitish grey, dense, sandy marl in the upper part. BIOSTRATIGRAPHY Nannoplankton In the Novouzensk key borehole, there are eight zonal complexes (Fig. 2) according to the standard nannoplakton zones of Martini (1971). Marl of the Algay Formation (Fm) (interval 926–917 m) contains an assemblage that is assigned to nannoplankton zone NP2 Cruciplacilithus tenuis, and includes C. tenuis, C. primus, M. inversus and rare specimens of B. sparsus. Zone NP3 Chiasmolithus danicus is recognised in the upper part of Algay Fm (917–892 m) and includes C. danicus, C. tenuis and rare specimens of B. sparsus. The Cygan Fm (892–856 m) is assigned to nannoplankton zone NP4 Coccolithus robustus (CP3 Ellipsolithus macellus according to Okada and Bukry 1980). Species present include C. robustus, C. eopelagicus, C. cavus,

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The Upper Syzran Subformation (789–721 m) contains nannoplankton of Zone NP5 Fasciculithus tympaniformis. This association includes F. tympaniformis, F. magnus, F. involutus, F. billii, F. janii, C. robustus, C. tenuis, C. primus, C. danicus and N. junctus. Nannofossils characteristic of Zone NP6 Heliolithus kleinpelli occur in the lower part of the Novouzensk Fm (721–696 m), where the first occurrence of H. kleinpelli is recorded. Rare specimens of Fasciculithus tympaniformis, C. tenuis, C. danicus are also present in this interval. The middle and upper parts of the Novouzensk Fm, along with the Bostandyk beds, lack nannoplankton species. The middle part of the Bostandyk Fm (475– 420 m) yields nannoplankton of the Zone NP12 Marthasteritus tribrachiatus. The upper Bostandyk Fm (420–414 m) contains abundant nannofossils and rich assemblages (ca 60 species), and is assigned to Zone NP13 Discoaster lodoensis. It includes Discoaster lodoensis, Braarudosphaera bigelowii, large specimens of Imperiaster obscurus, Discoaster kuepperi and Chiphragmalithus armatus. The top of the Bostandyk Fm (414–405 m) are noncalcareous and nannofossils are not present. The top of the section (405–400 m), Kopterek Fm, contains nannoplankton of Zone NP14 Discoaster sublodoensis. The presence of zonal species Rhabdosphaera inflata denotes the upper part of NP14 and is dated as Lutetian. This complex is rich (70 species) and has a very similar composition coeval assemblages of southern Russia except that species of Rhabdosphaera, Pontosphaera, Micrantolithus are extremely diverse.

Extended Abstracts


Figure 2 borehole. Extended Abstracts

The lithological sequence, the nannoplankton and dinocysts zones in the Novouzensk key

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Dinoflagellate cysts Organic-walled phytoplankton (dinocysts and acritarchs) are consistently present throughout the Novouzensk section. Studied zonal assemblages correspond to the dinoflagellate cyst zonation scheme in Luterbacher et al. 2004. Dinocysts assemblages from the PETM are correlated with North Sea Basin dinocysts zones (Mudge and Bujak 1994, 2001). Eleven dinocysts biozones are defined for the upper Danian–Lutetian interval (Fig. 2). Pollen and spores are also present in almost every interval of the borehole. The Algay Formation has not been examined. The Cygan Fm (856–869 m) contains dinocyst Zone D3a Alterbidinium circulum, and includes Alterbidinium circulum, Spinidinium densispinatum, Palaeocystodinium australinum, P. bulliformum, Senegalinium iterlaaense, Cerodinium striatum and C. kangiliense. The Lower Syzran Subformation (789–856 m) and the lower part of the Upper Syzran Subformation are assigned to dinocyst Zone D3b Cerodinium depressum. The first appearances of Cerodinium depressum and abundant Spinidinium densispinatum occur at the base of this biozone. Typical species include Cerodinium speciosum, Senegalinium iterlaaense and Palaeoperidinium pyrophorum. The upper dinocyst Zone D3b (DP3b) Isabelidinium? viborgense is recognized in the Upper Syzran Subformation (772–721 m), based on comparison with the North Sea Basin zonation of Mudge and Bujak (2001). The first and the last appearances of I. ?viborgense occur in this interval, along with the last appearance of species Palaeoperidinium pyrophorum and Palaeocystodinium australinum. The lower part of the Novouzensk Fm (721– 635 m) contains dinocysts of a regional, mainly endemic, assemblage referred to as “Beds with Cerodinium markovae”. It includes the zonal taxon, С. leptodermum, and rare specimens of Areoligera gippingensis and Alisocysta margarita. This association corresponds to Zone D4b (Luterbacher et al. 2004) and Zone DP5a-b of Mudge and Bujak (2001). Pollen and spores are dominant in this interval. The upper part of the Novouzensk Fm

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(568–523 m) is assigned to dinocyst Zone D4c (DP6a) Apectodinium hyperacanthum. The assemblage contains Apectodinium homomorphum, A. hyperacanthum, Deflandrea oebisfeldensis, D. andromiensis, Cerodinium speciosum subsp. glabrum and C. sibiricum. The top of the Novouzensk Fm is assigned to Zone D5a (DP6b) Apectodinium augustum. Species recorded include A. augustum, A. parvum, A. homomorphum, A. hyperacanthum, A. quinqelatum and W. pechoricum. The species Deflandrea oebisfeldensis is also abundant in this zonal association. This dinocyst zone corresponds to PETM interval. The lower part of the Bostanyk sandstone beds (523–500 m) contain only a small number of dinocysts and acritarchs dominate the assemblage. The main acritarchs are Pterospermella spp. and Leiosphaeridia spp. Dinocyst species Pyxidinopsis densepunctata is present and D. oebisfeldensis, C. wardenense, C. dartmoorium and Apectodinium sp. are rare. This assemblage corresponds with Zone DE1a of the North Sea Basin zonation scheme (Mudge and Bujak 2001). Zone DE1b-c Deflandrea oebisfeldensis is present in the upper part of the Bostandyk beds (500– 486 m) and the zonal species is dominant in this interval. The middle part of the Bostanyk Fm include a diverse dinocyst assemblage that is assigned to Zone D7c Dracodinium varielongitudum (Luterbacher et al. 2004). This interval can also be correlated with Zone DE2 of the North Sea Basin zonation (Mudge and Bujak 1994, 2001). The upper part of the Bostandyk Fm (460–414 m) contains an abundant and very diverse dinocyst assemblage that is assigned to Zone D8 Charlesdowniea coleothrypta (Luterbacher et al. 2004). The first appearance of Dracodinium politum, Areoligera medusettiformis, Wetzeliella samlandica and Pentadinium laticinctum mark the base of this zone in this section. In addition, Ochetodinium romanum, Heslertonia heslertonense, Charlesdowniea coleothrypta (rare), C. aff. clathrata and Kallosphaeridium brevibarbatum are present. The top of the Bostandyk Fm and Kopterek Fm are a non-calcareous silt clay and marl that is assigned to dinocyst Zone D9a Areosphaeridium diktyoplokum (Luterbacher et

Extended Abstracts


al. 2004). The first appearance of the zonal species, Wilsonidium echinosuturatum and Wetzeliella coronata, occur at the base of this zone. Marl from the top of the section contains a similar dinocyst association, including Wetzeliella ovalis, W. aff. articulata and Deflandrea apiculiformis, and is tentatively correlated with Zone D9 (Luterbacher et al. 2004).

dinocyst assemblage with the North Sea Basin Zone DE1a Leiosphaeridia spp. and Zone DE1b-c Deflandrea oebisfeldensis. •

A complete Danian–lower Lutetian nannoplankton and dinocyst zonal succession is present in the Northern Pricaspian Depression from the Novouzensk key borehole. The age is determined by the correlation of nannoplankton and dinocysts biozones.

The middle part of the Bostandyk Fm include nannoplankton Zone NP12 Marthasterites tribrachiatus and Zone NP13 Discoaster lodoensis, and dinocyst Zone D7c Dracodinium varielongitudum and Zone D8 Charlesdowniea coleothrypta. The top of the Bostandyk Fm contains a dinocyst assemblage assigned to Zone D9a Areosphaeridium diktyoplokum. Hence, the Bostandyk Fm are dated as middle-upper Ypresian.

•

The Algay Fm is characterized by nannoplankton Zone NP2 Cruciplacolithus tenuis and NP3 Chiasmolithus danicus, and is dated as lower–middle Danian.

The Kopterek Fm is assigned to nannoplankton upper Zone NP14 Discoaster sublodoensis and dinocyst Zone D9. These beds are dated as lower Lutetian.

•

Four hiatuses during the Danian–lower Lutetian are recognised in the Novouzensk section, based on micropaleontological analyses. The intervals of non-deposition (Fig. 2) appear to be (1) The base of the Lower Danian (Zone NP1); (2) Upper Selandian sediments are most possibly absent (top of the Upper Syzran Subformation), because the dinocyst Zone DP4a-b Palaeoperidinium pyrophorum is not present; (3) An Early Ypresian interval of non-deposition, including dinocyst Zone D6–D7a-b; and (4) A hiatus in the upper Ypresian–lower Lutetian that is younger than dinocyst Zone D9a Areosphaeridium diktyoplkum and continues into the upper part of nannoplankton Zone NP14 Discoaster sublodoensis.

CONCLUSIONS •

•

•

The Lower Syzran Subformation is represented by nannoplankton upper Zone NP4 Coccolithus robustus (Lone Neochiastozygus junctus), and also dinocyst Zone D3b Cerodinium depressum. The Danian–Selandian boundary is placed at the base of the Lower Syzran Subformation and coincides with a sediment change from calcareousterrigenous to siliceous-terrgeneous.

•

The Upper Syzran Subformation is assigned to nannoplakton Zone NP5 Fasciculithus tympaniformis and upper dinocyst Zone D3b Cerodinium depressum (DP3b Isabelidinium ?viborgense), dated as Selandian.

•

•

The Novouzensk Fm consists primarily of non-calcareous clay. The nannoplankton Zone NP6 Heliolithus riedelli is recognised at the base of this unit. The presence of the dinocyst species Cerodinium markovae in the lower part, and Zone D4c Apectodinium hyperacanthum in the upper part, indicate a Thanetian age for the Novouzensk Fm. The top of the Novouzensk Fm contains dinocysts zone D5a Apectodinium augustum and correlates to the PETM.

REFERENCES Luterbacher, H.P.; Ali, J.R.; Brinkhuis, H.; Gradstein, F.M.; Hooker, J.J.; Monechi, S.; Ogg, J.G.; Powell, J.; Rohl, U.; Sanfilippo, A.; Schmitz B. 2004: The Paleogene Period. In Gradstein, F.M.; Ogg, J.G.; Smith, A.G. ed. A Geological Time Scale. Cambridge, Cambridge University Press. Pp. 384-408. Martini E. 1971: Standard Tertiary and Quaternary calcareous nannoplankton zonation. In Farinacci, A. ed. Proceedings of the Second Planktonic Conference, Roma, 1970. Roma, Tecnoscenza. Pp. 739-785.

The Bostandyk Fm is dated as Early Ypresian, based on correlation of the

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Mudge D.C.; Bujak J.P. 1994: Eocene stratigraphy of the North Sea basin. Marine and Petroleum Geology 11: 166-181. Mudge D.C.; Bujak J.P. 2001: Biostratigraphic evidence for evolving Paleoenvironments in the Lower Paleogene of the Faeroe–Shetland

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Basin. Marine and Petroleum Geology 18: 577-590. Okada H.; Bukry D. 1980: Supplementary modification and introduction of code numbers to the low-latitude coccolith biostratigraphie zonation (Bukry, 1973, 1975). Marine Micropaleontology 5: 321-325.

Extended Abstracts


GEOCHEMISTRY OF “GLASSY” PLANKTONIC FORAMINIFERA FROM THE UPPERMOST EOCENE: IMPLICATIONS OF VITAL EFFECTS FOR RECONSTRUCTING SEA SURFACE TEMPERATURES Bridget S. Wade1,3, Andrea Dutton2, Yair Rosenthal3, Richard K. Olsson4, William A. Berggren4 and James D. Wright4 1

M.T. Halbouty Building, Department of Geology and Geophysics, Texas A&M University, College Station, TX 77843-3115, USA: [email protected]; 2Research School of Earth Sciences, Australian National University, Canberra, Australia; 3Institute of Marine and Coastal Science, Rutgers University, New Jersey, USA; 4 Department of Earth and Planetary Sciences, Rutgers University, New Jersey, USA.

Planktonic foraminifera are extensively used in paleoceanographic studies; however, many modern species of planktonic foraminifera do not record geochemical signals that reflect sea surface conditions. The reliability of planktonic foraminiferal geochemistry for documenting climate change is influenced by a number of factors including ontogenetic changes in element concentrations and stable isotopes (Erez 2003; Weiner and Dove 2003), diagenesis (Pearson et al. 2007), the presence of photosymbionts (Wade et al. 2008) and position within the water column. Therefore, documenting biological and ecological effects on geochemical proxies within different species of extinct planktonic foraminifera is necessary for determining the reliability of each species as recorders of paleoceanographic change. The latest Eocene is a critical interval of climatic change (Miller et al. 2008; Katz et al. 2008; Pearson et al. 2008); documenting the depth habitats of upper Eocene planktonic foraminifera allows the structure of the water column and its stability to be examined.

species, we conducted geochemical analysis on planktonic foraminifera from the Shubuta Formation, Mississippi, USA (Fig. 1). The planktonic foraminiferal assemblage is of Late Eocene age, Biochron E16, ca 34 Ma. While the bio- and sequence-stratigraphy of the US Gulf Coast region have been extensively studied (e.g. Keller et al. 1985; Mancini and Waters, 1986; Fluegmann et al. 1996; Miller et al. 2008 and references therein), the geochemistry of planktonic foraminifera has not been examined. The objectives of this study were to examine multiple species from multiple size fractions to: •

determine the paleobiology of extinct planktonic foraminifera;

•

document Late Eocene sea surface temperatures and the structure of the water column; and

•

examine intraspecific variations in geochemical signals through ontogeny.

To determine the paleobiology of extinct

Figure 1 Location of the study area (Shubuta type section, Wayne County, Mississippi). Extended Abstracts

Planktonic foraminifera are well preserved and comparable to assemblages from this interval from Tanzania (Wade and Pearson 2008). Specimens appear glassy under the light microscope with excellent preservation confirmed with the scanning electron microscope. Stable isotopic analysis of planktonic foraminifera was performed at Rutgers University. Foraminifera were reacted in phosphoric acid at 90°C for 15 minutes in an automated peripheral attached to a Micromass Optima mass spectrometer. Analytical precision measured using NBS-19 was 0.08 for δ18O and 0.05 for δ13C. 155


Standard procedures were used for cleaning planktonic foraminifera for trace element ratios. We included physical withdrawal of contaminants using the binocular microscope, a clay-removal step, two oxidation steps, a reductive step and one dilute-acid leach. Foraminiferal Mg/Ca ratios were measured using high-resolution inductively-coupledplasma mass spectrometry at the Institute of Marine and Coastal Sciences at Rutgers University (long-term precision 35 km of badland exposures in the southeastern BHB, measured local lithological sections, described paleosols, collected fossils, and measured the δ13C of various source materials. We attained stratigraphic precision of ±1 m within local areas using a jake staff and sighting level, and by taking coordinates with a differential GPS. We correlated local sections by recording bed traces with the GPS and reconstructing dip and strike from these measurements. We checked correlations with stable isotope and biostratigraphy, yielding a detailed stratigraphic framework for the study area.

(Fig. 2) and/or the first occurrence of the distinctive Wa0 mammalian fauna, though the latter is slightly later (Gingerich 2003). The base is marked lithologically by the first laterally persistent red paleosol, which typically contains small, pedogenic CaCO3 nodules (2 cm), irregular CaCO3 nodules and burrow fills. Nodule lags in channel deposits are common. Abandoned channel deposits 2–3 m deep and filled with mud or interlaminated silt and sand are common, and some preserve plant fossils.

I. Transitional Sequence

IV. Big Red Sequence

The uppermost 20–40 m of Fort Union Fm is similar to rocks below: drab mudstone paleosols, lenticular and tabular channel sandstones, and laterally extensive carbonaceous shales. The highest carbonaceous shale is 13 m below the base of the PETM. The Transitional Sequence differs from underlying Fort Union strata in having laterally discontinuous, reddish mudstone paleosols, suggesting locally better-drained floodplains. Near the eastern margin of the basin these reddish paleosols are more common and occur lower stratigraphically. In most of the study area the highest Fort Union unit is a laterally extensive fine to mediumgrained, poorly-sorted sandstone from 2–10 m thick that weathers in a steep face. This tabular body is locally downcut 10 m of underlying paleosols, leaving only the lower part of the Big Red Sequence. The appearance of the Big Red Sequence is also influenced by the thickness of the intercalated gray-white siltstone/sandstone units. At the east margin of the basin the three lowest paleosols in the Big Red Sequence appear to merge in weathered outcrop, forming a single red bed more than 5 m thick.

LITHOSTRATIGRAPHY

II. Lower Red Sequence We recognize the base of the PETM by a negative excursion in carbon isotope values 158

Extended Abstracts


Figure 2 PETM sections in the southeastern Bighorn Basin showing lithology, key fossil localities, and isotope stratigraphy. Colours in lithologic columns approximate paleosol colors in weathered outcrop.

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PALEONTOLOGY The Transitional Sequence has abundant fossil plants in both carbonaceous shales (wet floodplain habitats) and interlaminated siltstone/sandstones beds (near-channel habitats). The highest plants in the Fort Union Fm are 8 m below the first laterally extensive red paleosol and 18 m below the lowest Wa0 mammals in the Picnic Area Hill section (Fig. 2). Uppermost Fort Union floras are dominated by temperate deciduous groups such as Betulaceae, Platanaceae, Cercidiphyllaceae, Fagaceae, Sapindaceae (Aesculus), Ginkgo, and taxodiaceous conifers (Metasequoia and Glyptostrobus). Probable evergreen taxa include Lauraceae, palms and cycads. Most of the deciduous genera are common throughout the Late Paleocene and across the Holarctic region (Collinson and Hooker 2003). Palynofloras are also typically Late Paleocene, dominated by Taxodiaceae, Juglandaceae, and Betulaceae. Although leaf and pollen assemblages of Sequence I are compositionally similar to other Late Paleocene floras, the absence of some common leaf types with toothed margins causes leaf margin estimates of mean annual temperature to be higher than for any other Paleocene level in the BHB (Wing and Currano 2008). Vertebrate fossils are rare in the uppermost Fort Union Fm of the study area, with the exception of an ironcemented, sand-granule conglomerate that outcrops locally at the base of the tabular sandstone at the top of the formation. A vertebrate locality ~6 m below the base of the CIE has produced >200 specimens. The fauna includes champsosaurs, Aletodon gunnelli, Apheliscus nitidus, Haplomylus simpsoni, and Copecion brachypternus, lacks Plesiadapis cookei, and has a high relative abundance of Phenacodus and Ectocion. The presence of Copecion and absence of characteristic Eocene species indicates a latest Clarkforkian age (Cf3) (Secord et al. 2006). An in situ jaw of the Clarkforkian species Coryphodon proterus was found just 2.5 m below the base of the tabular sand at Big Red Spit (Fig. 2). Yans et al. (2006) reported a single specimen of the PETM index mammal, Meniscotherium, 3–4 m below the “gray-tracer” bed, which we have traced into a laterally extensive red paleosol that marks the base of the CIE in four other sections. We think this specimen is likely to have washed down slope, a common

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occurrence in this area because the steep-faced tabular sandstone underlying the base of the PETM facilitates downslope movement of abundant Wa0 vertebrate fossils into the sparsely fossiliferous Fort Union Fm below. The Lower Red Sequence (II), representing the first part of the PETM, has produced plant fossils in only two locations. Both are minor siltstone channel-fills 1–3 m thick and