Antarctic Climate Evolution, Volume 8 (Developments in Earth and Environmental Sciences)

Developments in Earth & Environmental Sciences, 8 ANTARCTIC CLIMATE EVOLUTION Edited by FABIO FLORINDO Istituto Nazio...

Author: Fabio Florindo | Martin Siegert

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Developments in Earth & Environmental Sciences, 8

ANTARCTIC CLIMATE EVOLUTION Edited by

FABIO FLORINDO Istituto Nazionale di Geoﬁsica e Vulcanologia, 00143 Roma, Italy

MARTIN SIEGERT School of GeoSciences, Grant Institute, University of Edinburgh, Edinburgh EH9 3JW, UK

Amsterdam – Boston – Heidelberg – London – New York – Oxford Paris – San Diego – San Francisco – Singapore – Sydney – Tokyo

Preface Antarctic Climate Evolution is the ﬁrst book dedicated to understanding the history of the world’s largest ice sheet and, in particular, how it responded to and inﬂuenced climate change during the Cenozoic. To explain the story of Antarctic ice and climate history, information on terrestrial and marine geology, sedimentology, glacier geophysics, ship-borne geophysics, and numerical ice sheet and climate modelling is presented within thirteen chapters. The book’s content largely mirrors the structure of the Antarctic Climate Evolution (ACE) program (www.ace.scar.org), an international initiative of the Scientiﬁc Committee on Antarctic Research (SCAR), afﬁliated with the International Polar Year 2007–2009, to investigate past changes in Antarctica by linking climate and ice sheet modelling studies with terrestrial and marine geological and geophysical evidence of past changes. The programme is designed to determine climate conditions and change in both the recent past (i.e. during the last glacial maximum, when temperatures were cooler than at present) and the more distant past (i.e. in the pre-Quaternary, when global temperature was several degrees higher than it is today). This new cross-disciplinary approach has led to a substantial improvement in the knowledge base on past Antarctic climate and to our understanding of the factors that have guided its evolution. This in turn has allowed us to build hypotheses, examinable through numerical modelling, for how the Antarctic climate is likely to respond to present and future global changes. Most of the subcommittees in ACE have been responsible for individual chapters, and in this way we have been able to cover the complete history of the Antarctic Ice Sheet and its climate evolution. The book will be of interest to research scientists from a wide range of disciplines including glaciology, palaeoclimatology, sedimentology, climate change, environmental science, oceanography and palaeoentology. It will also be valuable as a supplementary text for undergraduate courses. We are grateful to our many friends and colleagues for advice and encouragement through the gestation of the book over the last 3 years. We also acknowledge input to the ACE initiative by a number of scientists (many of them contributed to this book), including P. Barrett, A.K. Cooper, J. Francis, R. Gersonde, M.J. Hambrey, D.H. Harwood, A. Moldonado, D. Pollard,

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D. Sugden, G. Villa, P.-N. Webb and G.S. Wilson. We are sure that the chapter authors will join us in thanking the reviewers for their comprehensive and valuable comments and suggestions. We acknowledge their very special contributions to this book by naming them here: J. Evans, J. Francis, W. Howard, L. Krissek, A. Mackintosh, C. O’Cofaigh, G. Orombelli, A.H. Orsi, D. Pollard, C.A. Ricci, I.C. Rutt, E. Stump, C. Summerhayes and G.S. Wilson. Finally we thank Linda Versteeg-buschman, Femke Wallien and Suja Narayana of Elsevier Science for their support in the production of this book.

Fabio Florindo Martin Siegert Rome and Edinburgh, July 2008

Elsevier Radarweg 29, PO Box 211, 1000 AE Amsterdam, The Netherlands The Boulevard, Langford Lane, Kidlington, Oxford OX5 1GB, UK First edition 2009 Copyright r 2009 Elsevier B.V. All rights reserved No part of this publication may be reproduced, stored in a retrieval system or transmitted in any form or by any means electronic, mechanical, photocopying, recording or otherwise without the prior written permission of the publisher Permissions may be sought directly from Elsevier’s Science & Technology Rights Department in Oxford, UK: phone (þ44) (0) 1865 843830; fax (þ44) (0) 1865 853333; email: [email protected]. Alternatively you can submit your request online by visiting the Elsevier web site at http://www.elsevier.com/locate/permissions, and selecting Obtaining permission to use Elsevier material Notice No responsibility is assumed by the publisher for any injury and/or damage to persons or property as a matter of products liability, negligence or otherwise, or from any use or operation of any methods, products, instructions or ideas contained in the material herein. Because of rapid advances in the medical sciences, in particular, independent veriﬁcation of diagnoses and drug dosages should be made British Library Cataloguing-in-Publication Data A catalogue record for this book is available from the British Library Library of Congress Cataloging-in-Publication Data A catalog record for this book is available from the Library of Congress ISBN: 978-0-444-52847-6 ISSN: 1571-9197 For information on all Elsevier publications visit our website at books.elsevier.com

Printed and bound in Hungary 09 10 11 12 13 10 9 8 7 6 5 4 3 2 1

Developments in Earth & Environmental Sciences, 8 F. Florindo and M. Siegert (Editors) r 2009 Elsevier B.V. All rights reserved DOI 10.1016/S1571-9197(08)00001-3

Chapter 1

Antarctic Climate Evolution Martin J. Siegert1, and Fabio Florindo2 1

School of GeoSciences, Grant Institute, University of Edinburgh, The King’s Buildings, West Mains Road, Edinburgh EH9 3JW, UK 2 Istituto Nazionale di Geoﬁsica e Vulcanologia, via di Vigna Murata 605, 00143 Roma, Italy

ABSTRACT Central to the understanding of global environmental change is an appreciation of how the Antarctic Ice Sheet interacts with climate. To comprehend the processes involved one must look into the geological record for evidence of past changes. For several decades international efforts have been made to determine the glacial and climate history of Antarctica and the Southern Ocean. Much of this information derives from studies of sedimentary sequences drilled in and around the continent. In addition, there have been numerous terrestrial geological expeditions to the mountains exposed above the ice surface usually close to the margin of the ice sheet. Holistic interpretation of these data is now being made, and hypotheses on the size and timing of past changes in Antarctica are being developed. In 2004, the Scientiﬁc Committee on Antarctic Research (SCAR) commissioned a scientiﬁc research programme on Antarctic Climate Evolution (ACE) to quantify the glacial and climate history of Antarctica. This book is a result of that programme, and documents, for the ﬁrst time, the state of knowledge concerning the ice and climate evolution of the Antarctic continent and its surrounding seas through the Cenozoic era.

1.1. Introduction The Antarctic Ice Sheet has existed for approximately 35 million years, but it has ﬂuctuated considerably and has been one of the major driving forces for Corresponding author. Tel.: +44(0)131 650 7543; Fax: +44(0)131 668 3184;

E-mail: [email protected] (M.J. Siegert).

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Quat.

Pleistocene Pliocene 5.33 Ma Miocene 23.03 Ma Oligocene

PALAEOGENE

TERTIARY

CENOZOIC

NEOGENE

1.81 Ma

33.9 ± 0.1 Ma Eocene 55.8 ± 0.2 Ma Palaeocene 65.5 ± 0.3 Ma

Figure 1.1: Geological time periods during the Cenozoic era. Dates listed on the right hand side are taken from Gradstein et al. (2004).

changes in global sea level and climate throughout the Cenozoic (Fig. 1.1). The rates, size and frequencies of these ﬂuctuations have been the subjects of considerable debate. Determination of the scale and rapidity of the response of large ice masses and associated sea ice to climatic forcing is of vital importance, because ice-volume variations lead to: (1) changing global sea levels on a scale of tens of metres or more, and (2) alteration to the capacity of ice sheets and sea ice as major heat sinks/insulators. It is thus important to assess the stability of the cryosphere under a warming climate (IPCC, 2007), particularly as icecore records have yielded evidence of a strong correlation between CO2 concentrations in the atmosphere and palaeotemperatures (Fig. 1.2). This concern is justiﬁed when CO2 levels are compared with those of the past. Since Antarctica is a major driver of Earth’s climate and sea level, much effort has been expended in deriving models of its behaviour. Some of these models have been successfully evaluated against modern conditions. In 2004, modelling the past record of ice-sheet behaviour in response to changes in climate (inferred from ice cores, sedimentary facies and seismic data), palaeoceanographic conditions (inferred from palaeoecology and climate proxies in ocean sediments) and palaeogeography (as recorded in landscape evolution) was seen as a critical next step, and became the focus of the ACE programme. The ACE programme facilitates research in the broad area of Antarctic climate evolution. The programme links new geophysical surveys and

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Figure 1.2: Variation in the Earth’s temperature during the last 65 million years, based on reconstructions from deep-marine oxygen isotope records. Future atmospheric temperature scenarios are based on IPCC (2001). Greenhouse trace gas projections are shown at top of diagram. Given the worse-case scenario, planetary temperatures could increase in 100–300 years to a level where, according to our knowledge of previous Antarctic glaciations, ice cover on Antarctica could not be sustained. The representation of palaeotemperatures is adapted from Crowley and Kim (1995). geological studies on and around the Antarctic continent with ice-sheet and climate modelling experiments. The programme determines past climate conditions and changes in both the recent past (i.e. during the Holocene, prior to anthropogenic impacts as well as at the last glacial maximum, when temperatures were cooler than at present) and the more distant past (i.e. in the pre-Quaternary, when global temperature were several degrees warmer

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than they are today). This cross-disciplinary approach, involving climate and ice-sheet modellers, geologists and geophysicists, has led to a substantial improvement in the knowledge base on past Antarctic climate, and our understanding of the factors that have guided its evolution. This in turn allows us to build hypotheses, examinable through numerical modelling, for how the Antarctic climate is likely to respond to future global change.

1.2. Antarctic Glacial History As is discussed in Chapter 7, the East Antarctic plate formed a signiﬁcant component of the Gondwanaland super-continent during the Jurassic. Since 180 Ma, this continent broke up into what are recognised today as distinct continental landmasses with the repositioning of Antarctica at southern polar latitudes in the Early Cretaceous (ca. 120 Ma). In spite of its polar position, Antarctica is thought to have remained mostly ice-free, vegetated, and with mean annual temperatures above freezing until the latter half of the Cenozoic (around 34 million years ago, Fig. 1.1), whereupon the continent became subject to repeated phases of glaciation at a variety of temporal and spatial scales. The southern continent and its surrounding ocean basins have been the target of numerous scientiﬁc expeditions and several scientiﬁc drilling project efforts that have led to signiﬁcant advances in understanding of Cenozoic climate evolution, oceanography, and biota of the Antarctic continent and the Southern Ocean. The deep-ocean records document clearly the long-term cooling of climates over the past 50 million years and large-scale variability in the last 3–5 million years. They also show events that are either abrupt or brief (e.g. the Paleocene warming event with a duration of less than 1 million years; the Middle-Eocene Climatic Optimum, MECO at ca. 41 Ma), or are marked by a distinct shift in the rate at which long-term changes occur (i.e. middle-Miocene increased cooling trend). The explanation for these events include changes in atmospheric gas concentrations (e.g. carbon dioxide and methane), opening of gateways with enhanced ocean circulation, peaks in orbital forcing resulting from Croll–Milankovitch cyclicities, interactions with northern hemisphere glaciations and others. Scientiﬁc drilling on the Antarctic continental shelf and upper slope, to examine the direct record of glaciation, has been sparse and has had signiﬁcant problems with recovery (o20% in diamict) using current Integrated Ocean Drilling Programme (IODP) techniques. Consequently, the linkages between Antarctic

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continental shelf and deep-ocean basin records are not well established and the basic problem of ice-sheet history remains unsolved. Proxy measurements (particularly oxygen isotopes) provide general details, but initiation, growth and extent of the ice sheets still are debated. 1.2.1. Late Eocene-Early Oligocene Cooling The Eocene to early Oligocene (focus of Chapter 8) represent a time of global cooling which was marked by reorganisation of global ocean circulation patterns and signiﬁcant turnovers in the marine and terrestrial biota (e.g. Berggren and Prothero, 1992) that culminates in the development of the ﬁrst Antarctic Ice Sheet and an important expansion of Antarctic ice volume. Global deep-sea oxygen isotope records indicate that this long-term cooling trend was not monotonic, but that it was interrupted by a series of abrupt short-term (ca. 1 million years) excursions in d18O (Zachos et al., 2001). Among these, the Oi-1 cooling event (Miller et al., 1991) at 33.55 Ma marked one of the most signiﬁcant global climatic deteriorations in the Cenozoic in response to the appearance of the ﬁrst continent-wide glaciation in Antarctica (e.g. Zachos et al., 1996). Coupled GCM/ice-sheet modelling has already been used to show that the formation of the East Antarctic Ice Sheet was triggered by a combination of gradual pCO2 lowering coupled with iceclimate feedbacks and orbital-forcing-induced cooling, rather than by the cooling associated with the opening of circumpolar seaways during the earliest Oligocene (e.g. Kennett et al., 1974; DeConto and Pollard, 2003; Lawver and Gahagan, 2003).

1.2.2. Oligocene–Miocene Boundary Mi-1 Glaciation The Oligocene–Miocene boundary (Chapter 9) marks a signiﬁcant transition in the development of the Antarctic cryosphere, where small dynamic ice sheets of the late Oligocene rapidly expanded to continental scale in the early Miocene. The transition is recorded in benthic foraminiferal d18O records as a positive 1.0 per mil shift, representing the ﬁrst of the Miocene glaciations (Mi-1). The climatic signiﬁcance of this was ﬁrst outlined by Zachos et al. (2001) who recognised the coincidence of the Oligocene–Miocene boundary and the Mi-1 isotope excursion with an unusual coincidence of low eccentricity and low-amplitude variability in obliquity of the Earth’s orbit. Mg/Ca reconstructions imply little or no change in temperature and that the ice-volume increase was equivalent to 90 m of sea level lowering (assuming a

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Pleistocene calibration of 0.11m d18O per 10 m of sea level). Sediment cores recovered in the Western Ross Sea indicate orbital modulation of the ice sheet during the transition, and corroborate proxy ocean records (Naish et al., 2001). It is argued that the Mi-1 event occurred as a consequence of the aforementioned unique setup of orbital parameters during an interval of declining CO2 that led to a prolonged period of cold summer orbits, during which time a large ice sheet established. This was then followed by warmer polar summers and enhanced melt from increased eccentricity and highamplitude variability in obliquity in the early Miocene, allowing the recovery of vegetation on the Antarctic craton. 1.2.3. Middle-Late Miocene Cooling The middle-to-late Miocene period represents a time of signiﬁcant ice-sheet expansion in Antarctica (Chapter 10). Interpretations of deep-sea isotope records and observations from geologic data from around the world suggest that the middle-Miocene encompassed a change from a period of warm climatic optimum, approximately 17–15 million years ago, to the beginning of major cooling between ca. 14.5–13.5 million years ago, and the formation of a quasi-permanent ice sheet on East Antarctica (Lewis et al., 2007). One outstanding question revolves around whether this cooling led to an ice sheet in East Antarctica that remained stable and in existence to the present day or underwent large-scale ﬂuctuations. New seismic-stratigraphic data from the Ross Sea reveal at least ﬁve major intervals of ice shelf advance and retreat in the middle-Miocene. Much of this ice is sourced in West Antarctica, suggesting the presence of a large and dynamic ice sheet in a part of Antarctica that is conventionally thought to be of lesser importance at this time. One of the most vexing questions concerns the stability of Antarctic climate and ice during the late Miocene and it has been the subject of almost continuous debate for more than 20 years. A variety of indicators from the McMurdo Dry Valleys suggest the maintenance of stable, hyper-arid, colddesert conditions since 13 Ma. However, microfossil studies (mainly diatoms) in the Transantarctic Mountains, and sedimentological work within Antarctic fjords are suggestive of signiﬁcant climatic dynamism extending from the late Miocene through the Plio-Pleistocene. A degree of heterogeneity in climate response is expected considering the size and diverse landscapes of Antarctica. Yet the existing state of knowledge is sufﬁciently contradictory that the community has evolved into two camps when it comes to describing the ice sheet in pre-Quaternary time: the ‘stabilists’ and ‘dynamicists’ (Miller and Mabin, 1998).

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1.2.4. Pliocene Record The Pliocene Epoch (Chapter 10) is a critical time for understanding the nature of the Antarctic Ice Sheet as IPCC (2001, 2007) projections of global temperature rise suggest that we will reach Pliocene levels within the next hundred years (Fig. 1.2). Indirect evidence, such as sea level changes and ocean ﬂoor sediments, suggests that ice volumes were subject to cyclical variability. It is believed that, since Northern Hemisphere ice sheets were not fully developed, sea level changes were driven by ﬂuctuations of the Antarctic Ice Sheet. Many scientists believe that it was the relatively unstable West Antarctic Ice Sheet that was responsible for these changes, but the role of the much larger East Antarctic Ice Sheet remains controversial. Key to this argument is the timing of the transition of the East Antarctic Ice Sheet from a polythermal, dynamic condition to a predominantly cold-base and stable state. Two opposing and vigorously defended views prevail. The long-standing view is that the East Antarctic Ice Sheet became stable in mid-Miocene time; evidence of which is primarily from the longevity of the landscape and well-dated surfaces and ash deposits in the Dry Valleys region along the western border of the Ross Sea. Another controversial view is that terrestrial glacial deposits, known as the Sirius Group, scattered in a number of locations through the Transantarctic Mountains, indicate dynamic ice-sheet conditions as recently as Pliocene time; based on diatom biostratigraphy and preserved vegetation. The latter viewpoint is supported by work on deposits known as the Pagodroma Group along the western side of the Lambert Glacier-Amery Ice Shelf drainage system. Each argument is internally consistent and the biggest challenge is to reconcile the differing views. If the East Antarctic Ice Sheet was indeed subject to major ﬂuctuations until Pliocene time then, taking into account IPCC projections, we have cause to be concerned about the possibility of the East Antarctic Ice Sheet reacting to climate change within the next few centuries.

1.2.5. Pleistocene Glacial Cycles (Intervals of Extreme Warmth and Cold) Studies of Antarctic ice cores show that Pleistocene climate variability in the different sectors of the southern high latitudes (detailed in Chapter 11) has occurred out of phase. This raises questions about the response of the southern high latitudes to external climate drivers, such as orbital insolation, solar variability and internal ampliﬁers such as thermohaline circulation and

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carbon cycle changes that operate at both Croll–Milankovitch and millennial– decadal time scales. These questions highlight a need for appropriate time series of climate variability from all sectors of the Southern Ocean. Recovery of sediment sequences with expanded Pleistocene sections, such as those from beneath the McMurdo Ice Shelf as undertaken by the ANDRILL programme (see www.andrill.org), permits the study of the structure and timing of glacial and interglacial cycles in the Southern Ocean at millennial time scales that extend well beyond the last four major climate cycles. In addition, several groups organised under the International Marine Global Change Study (IMAGES) programme have proposed to collect long piston cores for Pleistocene research from several sectors of the Southern Ocean. With new high-resolution Pleistocene time series from both the Antarctic margin and offshore sites, we can begin to determine if the abrupt climate changes that have been documented from the Atlantic and Indian sectors, and in polar ice cores, have also occurred in the Paciﬁc basin. During the last decade, many palaeoceanographic studies focused on millennial climate variability (e.g. on the Dansgaard-Oeschger (DO) events). They show that ocean thermohaline circulation is capable of becoming unstable in response to climate change and modiﬁcations to the cryosphere. The existing palaeoceanographic record documents mainly the North Atlantic Ocean, and modelling experiments have mostly explored the variability of North Atlantic Deep Water formation forced by fresh water ﬂux from ice surge events. However, Southern Ocean sea ice may be important during glacial periods, and the glacial ‘on/off’ modes of global circulation could also be linked to deep-water formation in the Southern Ocean. 1.2.6. Last Glacial Cycle and Deglaciation At the last glacial maximum (LGM, B21 ka), ice-sheet expansion in Antarctica was responsible for around 15 m of global sea level fall, with growth most likely taking place over continental shelves exposed as a result of sea level fall from the development of Northern Hemisphere ice sheets (Siegert, 2001). There are currently three different ideas about the onset of deglaciation: (1) changes in the water balance of the North Atlantic, the source region for much of the global thermohaline circulation, serve to propagate the deglacial signal worldwide; (2) changes in the Southern Ocean, as recorded in some ice cores, lead deglaciation as seen in Greenland ice; and (3) synchronicity in the timing of high-latitude climate change in both hemispheres, and some tropical records, suggest that tropical forcing is a key initiator of

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deglaciation. It may seem surprising that this controversy has not already been settled. The most important confound for establishing synchronicity, or its absence, among the available palaeoclimate records revolves around chronology development. It is notoriously difﬁcult to date LGM ice layers and sediments to an accuracy of better than 1–2 thousand years. It is also difﬁcult to separate local climate or geomorphic signals from large transformations that are regionally or globally important. What is needed to resolve the deglacial synchronicity issue are better records from rapidly deposited deglacial sequences across a range of longitudes and latitudes in the Southern Ocean that use sedimentary or glacial outlet indicators to directly track regional climate systems. Currently there are too few precisely dated records of the LGM from the Southern Ocean. Chapter 12 brings together several geological datasets relevant to the LGM in an attempt to constrain the ice volume at this important time of climate change.

1.3. Structure and Content of the Book Antarctic Climate Evolution presents the state of knowledge concerning the ice and climate history of the Antarctic continent and its surrounding seas throughout the Cenozoic. It begins with two chapters that provide the historical context in which palaeoclimate knowledge has been gathered. In Chapter 2, background to the International Polar Years (IPYs) is presented. These are periods in which major advances in our understanding of the Antarctic continent have occurred. Chapter 3 summarises the history of geological investigations on the continent. The next three chapters provide background information necessary in comprehending Cenozoic change in Antarctica. Chapter 4 discusses the role of the Southern Ocean in modulating and controlling ice and climate in Antarctica, Chapter 5 presents evidence of ice-sheet changes from studies of sea-ﬂoor sediments (obtained from drilling and seismic investigations) and Chapter 6 introduces the concept of numerical ice-sheet modelling as a powerful tool in quantifying former ice sheets. From this point, the book presents a series of chapters, each of which deals with a speciﬁc time period in Antarctic history, as detailed above. Our appreciation of Antarctic climate evolution is possible because of pioneering research undertaken over the past 100 or so years. It has involved thousands of academics from a variety of disciplines and, of course, nations. Major advances in our knowledge have come about as a result of large integrated programmes of activity. Organising logistics in Antarctica necessary to gather geological records often needs the ﬁnances and assistance

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of several nations. For example, offshore drilling has been supported by the IODP, and near-shore drilling has been organised through the ANDRILL consortium. Ice coring, subglacial lake exploration, numerical modelling and palaeoceanography each have multination programmes at their core (Chapter 13). All four IPYs (1882–1883, 1932–1933, 1957–1958 and 2007–2008) have offered momentum and encouragement for international multidisciplinary activities that have lasted well beyond their ofﬁcial time frames. For example, the third IPY (also known as the IGY) resulted in SCAR and the Antarctic Treaty. With several new ambitious programmes being either undertaken or planned within the fourth IPY, the next decade will undoubtedly see our knowledge of Antarctic Climate evolution develop considerably. Such knowledge will be critical to assessing how the Southern Ocean and ice sheets interrelate and feedback with global climate change.

REFERENCES Berggren, W. A., & Prothero, D. R. (1992). Eocene-oligocene climatic and biotic evolution: An overview. In: D. R. Prothero, & W. A. Berggren (Eds). EoceneOligocene Climatic and Biotic Evolution. Princeton University Press, Princeton, NJ, pp. 1–28. Crowley, T. J., & Kim, K.-Y. (1995). Comparison of longterm greenhouse projections with the geologic record. Geophy. Res. Lett., 22, 933–936. DeConto, R. M., & Pollard, D. (2003). Rapid Cenozoic glaciation of Antarctica triggered by declining atmospheric CO2. Nature, 421, 245–249. Gradstein, J. G., Ogg, A. G., Smith, F. P., Agterberg, W., & 36 others. (2004). A Geologic Time Scale 2004. Cambridge University Press, 589 pp. IPCC (2001). The Scientiﬁc Basis Contribution of Working Group I to the Third Assessment Report of the Intergovernmental Panel on Climate Change (IPCC). In: J. T. Houghton, Y. Ding, D. J. Griggs, M. Noguer, P. J. van der Linden, & D. Xiaosu (Eds). Cambridge University Press, Cambridge, UK, 944 pp. IPCC (2007). Climate Change 2007: The Physical Science Basis. Contribution of Working Group I to the Fourth Assessment Report of the Intergovernmental Panel on Climate Change. In: S. Solomon, D. Qin, M. Manning, Z. Chen, M. Marquis, K. B. Averyt, M. Tignor, & H. L. Miller (Eds). Cambridge University Press, Cambridge, UK, 996 pp. Kennett, J. P., Houtz, R. E., Andrews, P. B., Edwards, A. R., Gostin, V. A., Hajos, M., Hampton, M. A., Jenkins, D. G., Margolis, S. V., Ovenshine, A. T., & Perch-Nielsen, K. (1974). Development of the circum-Antarctic current. Science, 186, 144–147. Lawver, L. A., & Gahagan, L. M. (2003). Evolution of Cenozoic seaways in the circum-Antarctic region. Palaeogeogr., Palaeoclimatol., Palaeoecol., 198, 11–37.

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Lewis, A. R., Marchant, D. R., Ashworth, A. C., Hemming, S. R., & Machlus, M. L. (2007). Major middle Miocene global climate change: Evidence from East Antarctica and the Transantarctic mountains. Geol. Soc. Am. Bull., 119, 1449–1461, doi:10.1130B26134.1. Miller, K. G., Wright, J. D., & Fairbanks, R. G. (1991). Unlocking the ice house: Oligocene-Miocene oxygen isotopes, eustacy and margin erosion. J. Geophys. Res., 96, 6829–6848. Miller, M. F., & Mabin, M. C. G. (1998). Antarctic Neogene landscapes – in the refrigerator or in the deep freeze? GSA Today, 8, 1–2. Naish, T. R., Woolfe, K. J., Barrett, P. J., et al. (2001). Orbitally induced oscillations in the East Antarctic Ice Sheet at the Oligocene/Miocene Boundary. Nature, 413, 719–723. Siegert, M. J. (2001). Ice sheets and Late Quaternary environmental change. John Wiley and Sons Ltd., Chichester, UK, 231 pp. 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., Quinn, T. M., & Salamy, K. A. (1996). High resolution (104 years) deep sea foraminiferal stable isotope records of the Eocene-Oligocene climate transitions. Paleoceanography, 11, 251–266.

Developments in Earth & Environmental Sciences, 8 F. Florindo and M. Siegert (Editors) r 2009 Elsevier B.V. All rights reserved DOI 10.1016/S1571-9197(08)00002-5

Chapter 2

The International Polar Years: A History of Developments in Antarctic Climate Evolution Fabio Florindo1,, Antonio Meloni1 and Martin Siegert2 1

Istituto Nazionale di Geoﬁsica e Vulcanologia, via di Vigna Murata 605, 00143 Roma, Italy 2 School of GeoSciences, Grant Institute, University of Edinburgh, The King’s Buildings, West Mains Road, Edinburgh EH9 3JW, UK

ABSTRACT The ﬁrst three International Polar Years (IPYs; 1882–1883, 1932–1933, 1957– 1958) were major periods of intense multidisciplinary polar research, bringing signiﬁcant new insights into global processes and laying the foundation of knowledge of the polar regions for future decades. The fourth IPY (2007–2009) continues the tradition of international science years and is one of the most ambitious internationally coordinated scientiﬁc research programmes ever attempted. In contrast to the three previous IPYs, the new IPY incorporates research within social science and its interface with the natural sciences. The new IPY also includes a wide range of education and outreach activities, and a commitment to excite and train the next generation of polar researchers. We discuss brieﬂy the history of the IPYs, and their contribution to comprehending Antarctic Climate Evolution.

2.1. Introduction The polar regions play key roles in global climate change and have profoundly affected environments during the Cenozoic, inﬂuencing sea levels, atmospheric composition and dynamics, and ocean circulation. Starting from the end of the nineteenth century, several major internationally Corresponding author. Tel.: +39 0651860 383; Fax: +39 0651860 397;

E-mail: ﬂ[email protected] (F. Florindo).

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coordinated explorations of the polar regions have taken place, which have improved our understanding of them and how they inﬂuence the world. The most prominent periods of exploration were: the ﬁrst International Polar Year (IPY) of 1882–1883 (e.g. Heathcote and Armitage, 1959; Wood and Overland, 2006), the so-called Heroic Age of polar exploration, stimulated by the International Geographical Congress of 1895, which had made Antarctica a target, the second IPY (1932–1933) (e.g. Gerson 1958) and the International Geophysical Year (IGY) of 1957–1958 (e.g. Buedeler, 1957; Korsmo, 2007), later extended to include 1959 and which started life as the third IPY. These were major initiatives that involved an intense period of multidisciplinary polar research bringing signiﬁcant new insights into global processes and laying the foundation of knowledge of the polar regions for future decades. The ﬁrst two IPYs and the IGY involved an increasing number of countries and scientists, and produced unprecedented levels of knowledge and understanding in many ﬁelds of research. The 12 countries of the ﬁrst IPY grew to 67 in the IGY, in which some 5,000 scientists and support staff were engaged in Antarctica alone. They not only changed the way science was conducted in the polar regions, from single nation programmes to complex multinational collaborations, they also standardised measurements, made data freely available to all and initiated the system of World Data Centers. The fourth IPY, 2007–2009 (Allison et al., 2007), is one of the largest collaborative science programs ever attempted. It continues the tradition of international science years, includes multidisciplinary research operating in both polar regions, and involves some 50,000 participants from 63 countries. A wide range of scientiﬁc problems will be addressed, including issues related to society. It differs from the three previous IPYs in that it includes all natural science disciplines – not just physics and geophysics, it incorporates the social sciences, and it includes a wide range of education and outreach activities aimed at attracting the next generation of polar scientists and engaging the attention of the public and policy makers. In this chapter, we discuss brieﬂy the history of the polar years, and their contribution to comprehending Antarctic Climate Evolution (ACE).

2.2. The First International Polar Year (1882–1883) In August 1874, Captain Karl Weyprecht (1838–1881) (Figs. 2.1 and 2.2) returned from an Arctic expedition, of which he was leader. The

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Figure 2.1: Carl Weyprecht (1838–1881) ideas initiated the ﬁrst IPY (photo courtesy: Archive, Alfred Wegener Institute for Polar and Marine Research).

Austro-Hungarian North Pole Expedition (1871–1874) aimed to explore the northwest of Nowaja Semlja, in the search for the Northeast Passage. During that Arctic expedition on the ice-strengthened schooner ‘Admiral Tegetthoff’, they discovered Franz Joseph Land (890 km from the North Pole) and gathered valuable information about the drift of icebergs and about meteorological and magnetic conditions in the Arctic. Although it was a successful expedition, it occurred to Weyprecht that single nation expeditions of this nature, often having geographical discovery as their primary goal, could only advance the frontiers of scientiﬁc knowledge to a limited extent. Knowing that answers to the fundamental questions of meteorology and geophysics were most likely to be found near the Earth’s poles, Weyprecht

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Figure 2.2: The cover page of the ‘Illustrierten Wiener Extrablattes’ (Viennese Illustrated Special Edition), 25 September 1874. The cover heralds the return of the leaders of the Austro-Hungarian North Pole Expeditions, Carl Weyprecht and Julius Payer. became an avid advocate of internationally coordinated exploration of the polar regions; his views were inﬂuential in the formation of the largest coordinated series of scientiﬁc expeditions taken in the polar regions during the nineteenth century, namely what is now known as the ﬁrst IPY of 1882–1883.

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During the 48th Meeting of ‘German Naturalists and Physicians’ in Graz (18 September 1875), Weyprecht gave a lecture about the ‘Basic principles of Arctic research’, in which he suggested establishing a network of ﬁxed Arctic observation stations (Baker, 1982). In 1879, during the second International Meteorological Conference in Rome, these ideas were presented together with those of Georg von Neumayer (1826–1909), ﬁrst Director of the German Hydrographic Ofﬁce in Hamburg. Here it was recommended to discuss the erection of a number of observatories in the Arctic and Antarctic for simultaneous hourly meteorological and magnetic observations around the poles. These ideas generated international interest and, during the ﬁrst International Polar Conference at the German Hydrographical Ofﬁce in Hamburg (1–5 October 1879), an organising body called the International Polar Commission, chaired by Neumayer, was established. This commission included Denmark, Norway, Russia, Sweden, Finland, Germany, AustriaHungary, the Netherlands, France, the United States and Great Britain, with the assistance of the new Dominion of Canada. During that conference, the ﬁrst IPY was planned for the biennium 1881–1882. The following year, during the second International Polar Conference in Bern (7–9 August 1880) this commission agreed to postpone the IPY and declared that it would be held in 1882–1883 to coincide with a transit of Venus across the face of the Sun, on 6 December 1882. In doing so simultaneous observations from different places on the globe could be made to calculate the astronomical unit (AU=nearly 150 million kilometres; the distance between the Earth and Sun). Between 1 and 8 August 1881, during the third International Polar Conference, held in St. Petersburg, and 5 months after Weyprecht’s death on 29th March 1881 in Michelstadt, the International Polar Commission outlined the details of the ﬁrst IPY, to last from 1 August 1882 until 1 September 1883 (Heathcote and Armitage, 1959). In total, 12 countries participated in the ﬁrst IPY resulting in 15 coordinated expeditions to the poles (13 to the Arctic, and 2 to peri-Antarctic islands). Fourteen research stations were established (Fig. 2.3) where researchers conducted experiments and gathered data (hourly records) over the course of the year that would greatly enhance the basis of then current knowledge of the Earth’s magnetic ﬁeld, surface weather conditions and astronomy. Two of these stations were in the Southern Hemisphere: Orange Bay at the southern tip of Tierra del Fuego (established by France) and Moltke-Hafen at Royal Bay, South Georgia (established by Germany). Another 34 permanent stations were located outside Polar territories (e.g. Shanghai, Rio de Janeiro, Bombay) bringing the number of stations participating in the IPY to 48. This ﬁrst IPY was primarily focused on

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Figure 2.3: North polar chart, showing Arctic research stations during the ﬁrst IPY. During the ﬁrst IPY, eleven nations established fourteen principal research stations across the polar regions. Twelve stations were in the Arctic and two stations were in the Antarctic region (map from the Scottish Geographical Magazine, Volume I, No. 12, 1885 and scanned by the University of Texas Libraries at http://www.lib.utexas.edu/).

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physics – especially meteorology, magnetism and auroral studies, rather than on interdisciplinary work, but its investigations did extend, though locally and in a limited fashion, to other ﬁelds like botany, geology and zoology. One of the most signiﬁcant results of the ﬁrst IPY was the mapping of the Aurora Borealis – known at the time as the Northern Lights – showing that it often occurs in an almost circular belt centred on the north magnetic pole. A huge amount of polar data for that epoch became available to scientists. In the manner of the times, the data analysis was largely conﬁned to the primary single disciplines that were the focus for the IPY – meteorology, geomagnetism, auroral observations. Unfortunately there was no central archive for the data and no organisation to facilitate exchange of information. As a result, no major attempts were made to synthesise the information obtained, although some scientists did manage to pull together data from others to support their own individual studies of geomagnetism and auroras. Nonetheless, the ﬁrst IPY is regarded as an epoch-making event in which a major step forward was made in environmental scientiﬁc knowledge of importance not just for the polar regions but for the Earth as a whole.

2.3. The Second International Polar Year (1932–1933) The success of the ﬁrst IPY stimulated an expanded effort 50 years later to hold a second IPY. During the 1920s, while conducting high-altitude weather balloon observations, scientists detected extremely strong winds at heights of 10–15 km above the surface of the Earth; these are known today as the ‘jet stream’. One such scientist was Johannes Georgi, a meteorologist of the Maritime Institute of Hamburg. During a meeting of the Deutsche Seewarte (1927), Georgi proposed to investigate this phenomenon with a coordinated international research effort that would commence on the 50th anniversary of the ﬁrst IPY. In June 1928, an informal organisational meeting was held in London to discuss plans for the event and a year later, in 1929, the International Meteorological Organization (IMO), the predecessor of World Meteorological Organization (WMO), endorsed the effort and formed a commission to undertake planning for the second IPY. In August 1930, a ﬁrst meeting of the International Polar Commission was held in Leningrad (now St. Petersburg) under the presidency of Dr. D. La Cour of Denmark in order to organise and integrate the total effort. Delegates of 10 nations were present: Canada, Denmark, Finland, France, Germany, Japan, Norway, UK, USSR and the USA. A further 16 nations expressed an interest in this

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initiative. The aims and the scientiﬁc programme of the second IPY were presented during the second meeting of the International Polar Commission held in Innsbruck, Austria, in September 1931. When the second IPY started on 1 August 1932, forty-nine nations participated and, despite considerable economic challenges, due to the worldwide economic depression that began in 1929, it heralded advances in many ﬁelds: meteorology, atmospheric science, the mapping of the jet streams, ionospheric soundings and their role in radio communications, magnetic observations and, to a lesser extent, atmospheric electricity (Fig. 2.4). The IPY was notable for the ﬁrst massive deployment of the new and somewhat experimental radiosondes for upper atmosphere measurement. Ninety-four research stations were maintained in the Arctic during the second IPY and many of these are still active today. This was the time of the Great Depression, so funds were limited. Plans for the Antarctic suffered as a result, and were not pursued as originally planned. Chile established a station at Punta Arenas and Argentina in the South Orkneys. Meteorological observations were made by Norwegian whalers in the Southern Ocean. Magnetic observations were made in several locations in the Southern Hemisphere, including Christchurch (New Zealand), Watheroo (Australia) and Cape Town (South Africa).

Figure 2.4: Stuart McVeigh, member of the Canadian team at Chesterﬁeld Inlet on Hudson Bay, north-eastern Canada, holding airborne kite with a meteograph on it (photo courtesy: Department of Physics fonds, University of Saskatchewan Archives).

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Among the many exceptional men that participated in the second IPY, one ﬁgure stands out as symbolic – Rear-Admiral Richard E. Byrd of the USA (Fig. 2.5). Byrd had intended to sail for Antarctica in 1932 in the middle of the IPY, but due to the funding problems created by the Great Depression had to delay sailing until late 1933 and did not reach the Ross Sea until January 1934, so in one sense his expedition was not strictly speaking part of the IPY. It was noted for its exploitation of major advances in technology, in particular aviation, navigation, motor transport and radio. These technologies changed the nature of polar exploration from ship-borne or land-based expeditions using dogs and sledges, to parties employing machines and aircraft. Clearly, while those early years of ﬂying were adventurous, ﬂying in Antarctica entailed considerable danger. Nonetheless, the potential of utilising aircraft for exploration purposes was

Figure 2.5: Admiral Richard E. Byrd, ca. 1930 (photo courtesy: The Ohio State University Libraries; http://library.osu.edu/sites/exhibits/byrdﬂight/).

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huge, especially in Antarctica where so much remained undiscovered. Admiral Byrd knew this, having ﬂown over the South Pole during his ﬁrst Antarctic expedition in 1929, as well as over the North Pole. It is because he took full advantage of developments in aviation that he is now remembered as one of the most famous and inﬂuential polar explorers. During Byrd’s polar expeditions during the second IPY he spent ﬁve months wintering alone whilst operating a meteorological station – the ﬁrst-ever research station located inland from the coastal margin (Bolling Advance Base), on the Ross Ice Shelf, at the southern end of Roosevelt Island (123 miles from the sea). Just as in the ﬁrst IPY, research during the second IPY was not limited to the polar regions. Even so, although many observatories were located in equatorial regions, those responsible at the time did not change the name of the enterprise as the main emphasis was clearly on the polar regions. The global network of observations allowed, for the ﬁrst time, an appreciation of geophysical phenomena at a planetary level. A vast amount of data produced during the second IPY was collected and the Commission for the Polar Year established an ofﬁcial repository for IPY at the Danish Meteorological Institute in Copenhagen (Denmark). The second IPY ofﬁcially continued until 1 September 1933, just prior to Byrd’s departure from the USA.

2.4. The Third International Polar Year/International Geophysical Year (1957–1958) Following the success of the ﬁrst two polar years it seemed reasonable to agree that these events should occur every 50 years. However, so many scientiﬁc and technological improvements were made in a short time after the second IPY that already at the beginning of the ﬁfties many scientists believed a third coordinated IPY would allow major advances in our knowledge of Antarctica, which should not wait until the 50th anniversary of the second IPY. On 5 April 1950, in Silver Spring, Maryland, a small group of eminent physicists gathered to meet in an informal meeting in Van Allen’s home. Among these were the house owner James Van Allen (1914–2006), Lloyd Viel Berkner (1905–1967), Siegfried Frederick Singer (born in 1924) and Sidney Chapman (1888–1970) (Fig. 2.6), all of whom had been involved in research for military applications during the World War II. They realised the potential of the new technologies such as rockets, radar and numerous other geophysical techniques perfected during the war, and hoped to

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Figure 2.6: Sidney Chapman (1888–1970) was the President of CSAGI and guided much of the IGY planning. He and Lloyd Berkner were called the ‘fathers of the IGY’ (NOAA 200th Anniversary Celebration Website, http://celebrating200years.noaa.gov/). redirect these toward creating a basic knowledge of Antarctica and the world. This desire reﬂected the lack of focus on the Antarctic in the previous IPYs. In the months following the ‘Van Allen’s dinner’, other geophysicists joined the proponents and, during the summer of 1950, they presented their idea for a third IPY at the Conference on the Physics of the Ionosphere, at Pennsylvania State University. From this point onwards the plan moved to an international scale and a proposal was presented to the International Council of Scientiﬁc Unions (ICSU) that the time was right for a new IPY, which would fall in 1957–1958 and coincide with an expected sunspot maximum (from the standpoint of solar-terrestrial research, the period of the previous IPY was not a particularly interesting time because it was coincident with a minimum in the 11-year solar cycle!). ICSU endorsed the proposal and broadened its scope to include studies of the whole planet, rather than just polar studies. The program was renamed the IGY and ICSU established a special committee (Comite´ Spe´cial de l’Anne´e Ge´ophysique Internationale, CSAGI), headed by S. Chapman (president) (Fig. 2.6) and

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Figure 2.7: The ofﬁcial logo of the 1957–1958 International Geophysical Year. L.V. Berkner (vice-president), to act as the governing body for all IGY activities (Fig. 2.7). Care was taken to ensure that this committee would remain non-nationalistic, apolitical and geared towards a scientiﬁc agenda. The ﬁrst meeting of CSAGI took place in Brussels, Belgium, in July 1953, and was followed by a series of general assemblies and regional conferences (held for the Arctic, Antarctic, the Americas, Eastern Europe, Africa and the western Paciﬁc). At the general assemblies the main criteria for the IGY were established. The Soviet Union was not initially among the 26 proposing countries, probably as a result of the very peculiar cold war politics of the time. Nonetheless, how could a program aiming to be named ‘International’ take place without the participation of the Soviet Union and its numerous allies? Fortunately, a short while later also the Soviet ‘block’ joined the program with a very comprehensive scientiﬁc approach. This truly international effort, against a background of cold war mistrust and weapons escalation, was a remarkable achievement. The launch of the ﬁrst artiﬁcial satellite, ‘Sputnik 1’, on 4th of October 1957, gave a temporary political victory to the Soviet Union in the race for space; but the Americans were not far behind, launching Explorer 1 on 31 January 1958. These launches began the exponential growth in geophysical knowledge about the state of the planet that was to come from the use of space probes; it was the ﬁrst dramatic step towards the space-based remote sensing techniques so much in use nowadays. The IGY, began on 1 July 1957 and was completed on 31 December 1958, although a one year extension – the International Geophysical Cooperation – was allowed. With more than 10,000 scientists and 67 countries involved, the IGY was a signiﬁcant event in the history of science and surely the greatest international scientiﬁc enterprise in the middle of the twentieth century.

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Important discoveries were made in the ﬁelds of cosmic ray research, climatology, glaciology, oceanography, terrestrial atmosphere and the magnetic ﬁeld (Fig. 2.8). Explorer 1, in particular, brought the discovery of the so-called Van Allen radiation belts. Major deep counter currents were discovered in the ocean, for example beneath the Gulf Stream. Importantly for Antarctica, the IGY also witnessed major, internationally coordinated over-ice traverses. Using seismic exploration techniques developed several years earlier during a Norwegian-British-Swedish expedition to Dronning Maud Land (1947–1952), the ice volume of the Antarctic continent was estimated with a degree of accuracy for the ﬁrst time. It showed the continent to be overlain with a vast ice sheet, several kilometres thick in places, with a bed often depressed below sea level and a separation between East and West Antarctica. Several research stations were located in the continental interior, including a US base at South Pole, a Soviet base at the Magnetic South Pole, named Vostok Station, and a Soviet station near the Pole of Relative Inaccessibility, named Sovetskya. A plinth and bust of Lenin was placed at the exact site of the Pole of Relative Inaccessibility, facing towards Moscow; it remains there today and was visited for the ﬁrst time in nearly 50 years in 2006. The impact of the IGY on comprehending the dimensions of the Antarctic Ice Sheet cannot be overstated. The measurements taken remain, in some of the more remote regions, the only data ever collected on ice thickness. Perhaps the most important conclusion of the IGY was an appreciation that Antarctic scientiﬁc exploration was best served by international corporation, regardless of global geopolitics. Subsequent to the IPY, nations interested in Antarctic matters gathered to declare Antarctica free from commercial exploitation, and a site for scientiﬁc collaboration and data sharing. The Scientiﬁc Committee on Antarctic Research (SCAR) was established during the IGY from discussions that took place in 1957 and led to the ﬁrst meeting of SCAR held at the Hague, from 3–6 February 1958. Its objective was to develop scientiﬁc cooperation on research of continental scope, and to ensure that such research continued beyond the narrow conﬁnes of the IGY. As an apolitical independent organisation, and an integral part of ICSU, SCAR has served to integrate and facilitate research on Antarctic matters ever since. Its members represent the national academies of science in 34 countries (as of early 2008). The need for this kind of mechanism was evident to ICSU from an examination of the shortcomings of the ﬁrst and second IPYs, when the organisation underpinning of the activity and its follow-up was weak or non-existent, which meant that the data were less valuable than they might otherwise have been.

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Figure 2.8: ‘Poles’ was one of six posters (Earth, The Oceans, The Poles, Weather and Climate, Sun and Earth, Space) around which the National Academy of Sciences’ IGY Committee created its booklet ‘Planet Earth’. The numbers on this poster identify points discussed in the booklet (National Academy of Sciences, 1958) (courtesy: National Academy of Sciences).

For much the same reason, another innovation of the IGY was ICSU’s creation of the network of World Data Centers to be repositories for the large amounts of data collected by the participants. The emphasis on putting data into a centre so that it would be available to all was a novel departure.

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2.5. The Fourth International Polar Year (2007–2008) The fourth IPY marks the 50th anniversary of the IGY, and will be one of the most ambitious internationally coordinated scientiﬁc research programmes ever attempted. The ofﬁcial IPY 2007–2008 observing period will be from 1 March 2007 to 1 March 2009, in order to include two full annual seasonal cycles in the Arctic and in the Antarctic. Planning for the fourth IPY began at SCAR’s Shanghai meeting in 2002, and took off with support from ICSU in 2003 for an International Planning Group. For former IPYs, the rationale behind them lay in coordinating international activities to explore the unknown. While exploration of the geographical unknown is also an important aspect of the fourth IPY, the main emphasis now is on integrating international efforts so as to better comprehend the behaviour of the Earth system, and understand the roles of the poles in global change. This change in emphasis is readily apparent from the six IPY themes: status, change, global linkages, new frontiers, vantage point and the human dimension (see below). The fourth IPY involves over 200 projects endorsed by the ICSU/WMO Joint Committee that steers the process. The projects are either focused on the Arctic, or on the Antarctic, or are bipolar (Fig. 2.9). Just as impressive is the amount of endorsement and publicity that has surrounded the fourth IPY. For example, the Antarctic Treaty Parties, at their summit in Edinburgh on 19th June 2006, stated ‘We express our support for a successful IPY. We believe that the scientiﬁc research undertaken during the IPY will increase knowledge of the Antarctic and will yield a better understanding of the major terrestrial, ocean and atmospheric systems that control the planet. The polar regions are sensitive barometers of climate change, and we value their biodiversity. Their health is vital to the well-being of the Earth’s systems and its inhabitants’. As in the case of the third IPY relative to the second IPY, technology has advanced considerably in the interim, and we now have a network of remote sensing, navigational and communications satellites along with massive advances in information technology that together have changed the way in which science is done. Much can now be done remotely – e.g. automated weather stations reporting back to base via satellite. There is much less emphasis therefore in the fourth IPY than there was in the IGY on establishing bases. Instead much will be done by remote observation. Another new departure is that much of the data collected will form important contributions to computer-modelling investigations of polar processes, such as ice sheet changes. Given the huge quantities of data that

Figure 2.9: IPY planning chart as of 10 October 2007: the framework of hexagons provides a visual impression of how all the endorsed IPY projects are related in terms of geography and topic and how they may be linked (source: David Carlson, IPY 2007–2008, International Programme Ofﬁce, Cambridge, UK).

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are hoped to be gained, and the need to inform and run numerical models, the full results of the IPY may take several years to become clear. There will be even more emphasis in the fourth IPY than in the IGY on the capture, storage and retrieval of data for the beneﬁt of all. The polar regions are particularly important at this time because, among other things, they are much more sensitive than are other parts of the world to climate change, and they are experiencing signiﬁcant environmental change, which is having a profound impact on the ecosystems and human activity. Aside from that, through long-range climate connections what happens in the polar regions inﬂuences what happens elsewhere on the Earth. As for the past IPY/IGY initiatives, the fourth IPY will provide a framework to undertake projects that normally could not be achieved by any single nation. Thousands of scientists from more than 60 countries, including those not traditionally involved in polar research, will examine a wide range of physical, biological and social research topics. A departure from the three previous IPYs is that about 60 of the 230 or so projects will focus on education and outreach objectives, aimed at attracting the next generation of polar scientists and engaging the public in global environmental issues. As mentioned above, the various scientiﬁc proposals from the scientiﬁc community led the IPY organisers to identify six science themes, listed below. Some projects will contribute to more than one of these themes. The majority aim at understanding the changing polar environment and the impact of those changes (e.g. Allison et al., 2007): 1. Status: to determine the present environmental status of the polar regions; 2. Change: to quantify and understand past and present natural environmental and social changes in the poles and to improve projections of future change; 3. Global linkages: to advance understanding on all scales of the links and interactions between polar regions and the rest of the globe, and of the processes controlling these; 4. New frontiers: to investigate the frontiers of science in these regions; 5. Vantage point: to use the unique vantage point of the polar regions to develop and enhance observatories from the interior of the Earth to the Sun and the cosmos; 6. The human dimension: to investigate the cultural, historical and social processes that shape the sustainability of circumpolar human societies and to identify their unique contributions to global cultural diversity and citizenship.

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During the IPY there will be considerable emphasis on improving our understanding of the behaviour of the Antarctic Ice Sheet and the climate of the region, and on the history of the ice sheet and climate. The ACE programme will make a signiﬁcant contribution to this latter goal. IPY projects will use proxy records from sediment cores, ice cores and other sources to deﬁne how the past climate and environments changed. A greater understanding of past changes in this region is crucial to forming a better knowledge of future global environmental change and to predicting the role of the Antarctic Ice Sheet in the future as Earth warms. During the IPY new ice cores will be recovered in Greenland (North Greenland Eemian Ice Drilling, NEEM project), and Antarctica (West Antarctic Ice Sheet Divide, WAIS Divide), providing new high-resolution records of glacial and interglacial changes during the Quaternary. On a longer time scale, drilling of sediment cores along the Antarctic continental margin (ANDRILL; www.andrill.org) aims to study the glacial and climate history of Antarctica and the Southern Ocean following the late Cretaceous (e.g. see Florindo et al., 2003, for a review of the recent history of circum-Antarctic drilling by the Ocean Drilling Program and the Cape Roberts Project, and see Hambrey and Barrett, 1993, for a more comprehensive review of earlier drilling in the Ross Sea region). Finally, it should be mentioned that the IPY is not the only activity conducting special activities during the 50th anniversary of the IGY. There are also the International Year of Planet Earth (http://www.esfs.org/); the Electronic Geophysical Year (EGY) (http://www.egy.org/); the International Heliophysical Year (IHY) (http://ihy.gsfc.nasa.gov/). Various International Council for Science Unions are coordinating these efforts and the International Programme Ofﬁce, established by ICSU and WMO at the British Antarctic Survey in Cambridge, UK, will serve as the ofﬁcial contact point for these other programmes.

REFERENCES Allison, I., Be´land, M., Alverson, K., Bell, R., Carlson, D., Danell, K., Ellis-Evans, C., Fahrbach, E., Fanta, E., Fujii, Y., Glaser, G., Goldfarb, L., Hovelsrud, G., Huber, J., Kotlyakov, V., Krupnik, I., Lopez-Martinez, J., Mohr, T., Qin, D., Rachold, V., Rapley, C., Rogne, O., Sarukhanian, E., Summerhayes, C., & Xiao, C. (2007). The scope of science for the International Polar Year 2007–2008, WMO/TD, No.1364, pp. 1–79. Baker, F. (1982). The ﬁrst International Polar Year. Polar Record, 21 (132), 275–285.

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Buedeler, W. (1957). The International Geophysical Year. United Nations Educational, Scientiﬁc and Cultural Organisation (UNESCO), Oberthur Rennes, Paris, France, Vol. XV, pp. 1–72. Florindo, F., Cooper, A. K., & O’Brien, P. E. (2003). Antarctic Cenozoic palaeoenvironments: Geologic record and models. Palaeogeogr. Palaeoclimatol. Palaeoecol., 198(1–2), pp. 1–278. Gerson, N. C. (1958). From polar years to IGY. In: H. E. Landsberg, & J. Van Mieghem (Eds). Advances in Geophysics. Academic Press, New York, Vol. 5, pp. 1–52. Hambrey, M. J, Barrett, P. J. (1993). Cenozoic sedimentary and climatic record, Ross sea region, Antarctica. In: J. P. Kennett, & D. A. Warnke (Eds). The Antarctic Paleoenvironment: A Persepective on Global Change, Part 2. Antarctic Research Series. American Geophysical Union, pp. 91–124. Heathcote, N. de V., & Armitage, A. (1959). The First International Polar Year. In: Annals of the International Geophysical Year, Vol. 1. Pergamon Press: London, pp. 6–100. Korsmo, F. L. (2007). The genesis of the international geophysical year. Phys. Today, 60(7), 38–43. National Academies Archives (1958). IGY ‘‘Planet Earth’’ Poster and Booklet, http://www7.nationalacademies.org/archives/IGYPlanetEarthPosters.html, 1–44. Wood, K. R., & Overland, J. E. (2006). Climate lessons from the ﬁrst International Polar Year. Bull. Am. Meteorological Soc., 86, 1685–1697.

Developments in Earth & Environmental Sciences, 8 F. Florindo and M. Siegert (Editors) r 2009 Elsevier B.V. All rights reserved DOI 10.1016/S1571-9197(08)00003-7

Chapter 3

A History of Antarctic Cenozoic Glaciation – View from the Margin Peter Barrett Antarctic Research Centre, Victoria University of Wellington, P.O. Box 600, Wellington, New Zealand

ABSTRACT The scale and antiquity of the Antarctic Ice Sheet was sensed from the time of earliest exploration a century ago. However, signiﬁcant advances in scientiﬁc thinking, along with logistics and technology for gathering data from the continent itself, were required before a clear and consistent framework for ice-sheet history and behaviour could develop and this has emerged only in the last few years. The main features of the present ice sheet were established by over-snow traverses during and following the International Geophysical Year (1957–1958), but the timing and circumstances of its origin remained uncertain. Geological records of post-Jurassic time were largely buried under the ice or the sea ﬂoor around the Antarctic margin, though a few radiometric ages from the new K–Ar dating indicated Antarctic glaciation was likely older than the Northern Hemisphere ice ages of the Quaternary Period. New post-World War II techniques in offshore surveying with marine geophysics and ship-based drilling were ﬁrst applied to the Antarctic margin in the early 1970s, and were immediately productive. The Antarctic continental shelf was found to be underlain by sedimentary basins with the promise of ice-sheet history, and in early 1973, cores from Leg 28 of the Deep Sea Drilling Project (DSDP) in the Ross Sea provided the ﬁrst physical record of Antarctic glaciation extending back to Oligocene times. DSDP Leg 29 drilled in the Southern Ocean for deep-sea cores from the whole Cenozoic Era, yielding the ﬁrst set of oxygen isotopic ratios (d18O) from benthic calcareous microfossils, and the ﬁrst estimates of ice volume. These indicated a two-stage ice-sheet history (cooling and some ice

Corresponding author. Tel.: þ64 4 463 5336; Fax: þ64 4 463 6581;

E-mail: [email protected] (P. Barrett).

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at 34 Ma, and persistent ice from 14 Ma – time-scale of Berggren et al., 1995). About the same time, the Dry Valleys Drilling Project (DVDP) was launched to recover onshore records of Antarctic Cenozoic climate history, and exploratory drilling from fast ice offshore soon followed. Cores from the Ocean Drilling Program in the 1980s from the Prydz Bay Shelf and the Kerguelen Plateau established the timing of the ﬁrst continental ice sheet at 34 Ma. In the same period, deep continuous coring from sea ice in McMurdo Sound developed from DVDP technology succeeded in providing new detail for its subsequent history. Core recovery at B98% yielded lithological evidence of ice margin and sea level ﬂuctuations implied by deep-sea isotope records. Further drilling with improved chronology in the 1990s yielded cores conﬁrming ice margin and sea level changes on Milankovitch frequencies and on a scale of 10–40 m of sea level equivalent. Micropalaeontological and geochemical evidence pointed to a slight cooling from a coastal cold temperate climate, with beech forests during interglacial times. Subsequent development of ice-sheet modelling has indicated that most of the cooling that initiated ice-sheet glaciation was the consequence of a fall in atmospheric CO2 levels below a critical threshold, allowing ice sheets to form that were highly sensitive to orbital forcing. This claim has been supported by recent work on CO2 proxies and indicators of a shift in carbonate compensation depth in deep-sea sediments. Since the ﬁrst measurements in the 1970s, deep-sea isotopic measurements have implied a signiﬁcant increase in Antarctic ice volume at around 14 Ma that persisted to the present day. However, in the mid-1980s, marine diatoms in glacial deposits in the Transantarctic Mountains suggested periods in Pliocene times when seas invaded the East Antarctic interior, implying dynamic Antarctic Ice Sheets until Quaternary times. Evidence of continued cold in the mountains over the last 14 Ma, glaciological problems with the proposed over-riding scenario, lack of a signal in the deep-sea isotope record for the loss of most Antarctic ice in Pliocene times and possible alternative atmospheric sources for diatoms has shifted the weight of evidence in favour of persistent ice in the east Antarctic interior. However, coastal outcrops in Prydz Bay and a recent deep core from beneath the McMurdo Ice Shelf have shown that in the globally warmer Pliocene, notably around 3–5 Ma, seas around the Antarctic margin were several degrees warmer. Indeed, recent drill cores suggest that the Ross Embayment and perhaps also most of the West Antarctic interior were periodically ice-free in these times. Three decades ago, the Antarctic Ice Sheet was seen as a long-standing feature of the Earth with its origins in early Cenozoic times and its permanency assured by mid-Miocene cooling. Research in the last decade from geological drilling and glaciological remote sensing, supported by ice sheet and climate modelling, indicates the ice sheet is in fact quite responsive to changes in the global climate system, whether natural or human-induced, though at different rates in different sectors. Recent developments in both science and technology outlined here provide opportunities for projecting realistic scenarios for future ice-sheet response on human time scales.

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3.1. Introduction This chapter traverses the growth in knowledge and understanding of Antarctic Ice Sheet history through the Cenozoic Era over the last 100 years and brieﬂy considers the future. Important external inﬂuences on progress have been the advances in biostratigraphy and geochronology through the latter part of the last century, seismic stratigraphy and deep-sea isotopic studies world-wide over the last four decades, and the marked progress in modelling ice behaviour along with other components of the Earth system over the last two decades. Key events in this account are given in Table 3.1. Interest in the climate history of the Antarctic continent ﬁrst developed with the curiosity-driven scientiﬁc expeditions from the Northern Hemisphere beginning in the late nineteenth century (the ‘‘Heroic Era’’, Fogg, 1992, p. 108ff). These included prominent explorers and scientists from Belgium, Britain, France, Germany, Japan, Norway and the United States and within two decades the extent and salient features of the present ice sheet and the continent beneath had been documented. These included: (i) An ice sheet of around 5 million square miles (13 million square kilometers) in area and rising to an elevation of at least 10,000 ft (3,000 m), along with shelf and sea ice. In terms of its history, it was thought most likely to have been more extensive in warmer Pliocene times, and to have originated in Miocene times (Taylor, 1922), though Wright and Priestley (1922, Table 17) noted Antarctic glaciation ‘‘apparently began there in Eocene or Oligocene times’’; (ii) Flat-lying ‘‘cover beds’’ similar to those in the Transantarctic Mountains (Beacon Sandstone of Late Paleozoic age), with fossil plants similar to those in India, South Africa and South America, indicating a former temperate climate in those times (Seward, 1914); (iii) An East Antarctic ‘‘shield’’, with a foundation of Precambrian rocks and a faulted Ross Sea margin comparable to that of eastern Australia. West Antarctica was seen geologically as more related to southern South America and linked to it through the Scotia Arc (David, 1914). The decades that followed saw extensive inland exploration largely by the Byrd expeditions (Fogg, 1992, pp. 134–146), with detailed observations of weather, snow and ice cover and some geological observations, but there was little advance in comprehending the basic geological history of the continent. However, there were signiﬁcant advances in technology and logistics through the introduction of radio communication, seismic sounding, ships and aircraft. These advances prepared the way for US Operation Highjump

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Table 3.1: Events in progress in the evolution of antarctic glacial history. Date 1910–1912

1928–1947 1957–1964

1964–1970

1972

1970–1973

1973–1975

1978 1978–1980

1982 1984

Event Scale of Antarctic Ice Sheet recognized. Simple history proposed. Origin thought likely Miocene but possibly Eocene–Oligocene Inland exploration/technology development by Byrd expeditions International Geophysical Year (IGY) supports extensive geophysical exploration but still the lack of geological record between Jurassic and late Quaternary time. Glacial features mapped in McMurdo region thought Quaternary Miocene radiometric ages indicate pre-Quaternary glaciation a reality. Oligocene ice rafting in Southern Ocean Pre-Quaternary history thought likely from Northern Hemisphere record Recognition of widespread pre-Quaternary glacial deposits (Sirius Formation) on land in high Transantarctic Mountains Early seismic surveys of continental shelf by USNS Eltanin First drilling on shelf, showing Antarctic continental glaciation as old as 25 Ma (DSDP Leg 28) Drilling in deep-sea ﬂoor showing cooling and ﬁrst ice sheet at 34 Ma and present ice sheet dating from B14 Ma (DSDP Leg 29) First onshore scientiﬁc drilling (McMurdo Dry Valleys) leading to late Neogene glacial record going back at least 4 Ma in Taylor Valley. First attempt at drilling from ﬂoating ice (DVDP 15) Collapse of West Antarctic Ice Sheet projected from rising CO2 First hole drilled through the Ross Ice Shelf (J9). Sea ﬂoor cores recover early Miocene diamict with diatomite clasts and pollen Early simple Antarctic Ice Sheet model Marine diatoms in glacial deposits of high Transantarctic Mountains (Sirius Fm) suggest East Antarctic interior seas B3 Ma ago Results from ﬁrst extensive multichannel seismic survey of the Antarctic margin (R/V S.P. Lee, Ross Sea and Wilkes Land)

References Wright and Priestley (1922), Taylor (1914, 1922) Fogg (1992) Pe´we´ (1962), Bull et al. (1962), Nichols (1964)

Craddock et al. (1964), Margolis and Kennett (1970), Flint (1971)

Mercer (1972), Mayewski (1975) Houtz and Meijer (1970) Hayes, Frakes et al. (1975)

Kennett, Houtz et al. (1975) Smith (1981), Torii (1981)

Mercer (1978) Clough and Hansen (1979), Webb et al. (1979) Oerlemans (1982) Webb et al. (1984), Harwood (1986)

Cooper and Davey (1985), Eittrem and Hampton (1987)

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Table 3.1: (Continued ). Date 1986–1988

1987 1986–1996

1988–2002

1993

1997–1999

2000 2001

2003

2004

2005

2006–2007

Event New drilling results from Ross Sea (CIROS) and Prydz Bay (ODP Leg 119) show cyclicity in early history and ice-sheet antiquity Vostok ice core provides climate record through last glacial cycle Formation of the SCAR Group of Specialists on the Evolution of Cenozoic Palaeoenvironments of the Southern High Latitudes Formation of ANTOSTRAT – Antarctic Offshore Acoustic Stratigraphy to coordinate Antarctic margin studies Ice-sheet model and deep-sea isotope record inconsistent with major Pliocene deglaciation Pre-mid-Miocene age for Sirius deposits proposed on basis of ancient ash and persistent cold in high Transantarctic Mountains Cape Roberts cores yield high-resolution record of signiﬁcant orbitally forced ﬂuctuations in Antarctic Ice Sheet and sea level from 34 to 17 Ma and slight cooling over that period New drilling results from Prydz Bay (ODP Leg 188) Erice workshop reviews ANTOSTRAT and proposes new group to integrate geophysical/ geological data on Antarctic glacial history and the new generation of coupled ice– ocean–atmosphere models (ACE) Coupled ocean–atmosphere–ice-sheet model shows major role for CO2 in early ice-sheet formation EPICA yields Dome C climate record through last 8 glacial cycles Formation of ACE as a SCAR Scientiﬁc Research Program Landscape evolution model for Lambert drainage basin shows similar ﬂuvial and glacial inﬂuence ANDRILL results provide ﬁrst continuous record of late Antarctic Cenozoic history, including ﬁrst proximal record of midMiocene transition, and evidence of ice-free Ross embayment 3–4 Ma ago

References Barrett (1989), Barron, Larsen et al. (1989) Barnola et al. (1987) Webb (1990)

Cooper et al. (2002, this volume) Huybrechts (1993), Kennett and Hodell (1993) Sugden et al. (1993)

Naish et al. (2001), Barrett (2007)

O’Brein et al. (2001) Cooper et al. (2002)

DeConto and Pollard (2003) EPICA (2004) Siegert et al. (2004) Jamieson et al. (2005)

Naish et al. (2008a), Harwood et al. (2003)

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immediately following World War II. Although largely a polar training exercise, it provided much of the technological foundation for the International Geophysical Year (IGY) in Antarctica (1957–1958).

3.2. Mid-Twentieth Century Advances (1956–1972) IGY saw the establishment of a network of permanent bases around the continent from which to execute ship-based surveys of the margin, along with stations in the interior (Byrd, South Pole, Charcot and Vostok), enabling the ﬁrst great scientiﬁc over-snow traverses gathering geophysical and glaciological data across the continent (Fogg, 1992, pp. 168–176). The technical and scientiﬁc success of IGY led an agreement to extend the research that had been initiated and the establishment of a Special (later Scientiﬁc) Committee on Antarctic Research (SCAR) in 1958 to facilitate and help coordinate these activities. The success of IGY also led participating nations in the years immediately following to craft an Antarctic Treaty, which was ratiﬁed in 1961 (Beeby, 1972). Though geology was subordinate to the main mission of IGY, geologists enthusiastically took advantage of the extensive access to the continent made possible by their geophysical and glaciological colleagues. Geological surveys by contrast were carried out by smaller parties visiting the numerous but limited exposed patches of rock around the continent. By this time, there had been signiﬁcant advances in the Earth sciences since the Heroic Era, and the application of this new knowledge to the Earth’s only polar continent led to great interest in reports from the region. These advances included a robust geological time scale through radio-isotopic dating (Holmes, 1965), and the different character of the crust beneath oceans and continents recognized from seismology, along with the concept of continental drift from geology (Du Toit, 1937; Carey, 1958), later embodied in the theories of Sea Floor Spreading and Plate Tectonics (Dietz, 1961; Hess, 1962; Wilson, 1965). By 1970, the comprehensive surveys carried out in the IGY and the decade following had resulted in a comprehensive new view of the climate, physiography, glaciology, geology and biota of the Antarctic continent, summarized in the American Geographical Society’s map folio series (Bushnell, 1964–1975). Antarctic sedimentary strata overlying basement rocks were found to record a climate history similar to that of the other Gondwana continents, with dry warm conditions in Devonian times (390–340 Ma), ice-sheet glaciation in Carboniferous–early Permian times (340–280 Ma) and a humid temperate climate in Permian–Triassic times (280–200 Ma) with evidence of rivers, lakes and coal swamps. However, this long sedimentary record of deposition ended

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around 180 Ma ago with Jurassic igneous activity associated with the break-up of the Gondwana supercontinent, leaving a huge gap in the Antarctic geological record until the deposition of moraines around the Antarctic margin, then presumed to have formed from ice sheets in the Quaternary Period (Harrington, 1965). Early post-IGY surveys of glacial deposits in the McMurdo Sound region by Pewe (1960), Bull et al. (1962), Nichols (1964) and others described the features they found in terms of four glacial episodes within the Quaternary Period, perhaps inﬂuenced by the prevailing view prior to the IGY in the Northern Hemisphere (Flint, 1957) and a lack of suitable fossil material or radiometric techniques for a sound chronology. Indeed, Nichols (1964) in his review of the status of Antarctic glacial geology stated, ‘‘This writer believes that as yet there is no good evidence for Tertiary Antarctic glaciation’’. In the years that followed, K–Ar dating revealed a signiﬁcant preQuaternary history for Antarctic Cenozoic ice sheets. In the Jones Mountains in West Antarctica, a glaciated surface overlain by basalt was dated at more than 10 Ma (Craddock et al., 1964) and in Taylor Valley west of McMurdo Sound, small basaltic cones overlain and underlain by evidence of glaciation yielded ages ranging between 2.8 and 3.6 Ma (Armstrong et al., 1968; Denton et al., 1970). Post-IGY research world-wide was reporting evidence of pre-Quaternary glaciers in the Arctic and of late Cenozoic cooling from fossil molluscs and plants in lower latitudes, leading Flint (1971, p. 441) to conclude, ‘‘Glaciation occurred in the Miocene and Pliocene as well as in the Quaternary. Cold periods were more numerous than the four periods of the classical literature’’. Cores from the ﬂoor of the Southern Ocean suggested that Antarctic ice could have been much older, with icerafted sand grains as old as Oligocene (Margolis and Kennett, 1970). But there was no way of showing whether these came from local ice caps or continental ice sheets.

3.3. First Antarctic Drilling (1972–1975) The period following World War II was marked by an expansion in both the scientiﬁc exploration of the oceans and the search for oil on the world’s continental shelves. Echo-sounding techniques developed in war time were applied to marine seismic surveys for science and industry, and ship-based drilling for oil offshore attracted the attention of scientists for drilling through the Earth’s crust. By the mid-1960s, the main features and gross chronology of most of the world’s continental shelves were known by the oil

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industry, and the US marine science community persuaded its government to support the development of an international Deep Sea Drilling Project (DSDP) to drill the world’s ocean ﬂoor, beginning in 1968. The main aims were to test and reﬁne the new theory of Plate Tectonics as well as gain new knowledge from this unknown 70% of the Earth’s surface that might improve our knowledge of Earth history. DSDP was remarkably successful and in its ﬁrst 4 years had drilled at over 200 sites around the world but not in the Antarctic region. The ﬁrst venture south was Leg 28, designed to carry out the ﬁrst drilling transect between Australia and Antarctica for dating the timing of continental separation from the age of sediment resting on basaltic ocean ﬂoor, and to core for climate and ocean history in the Southern Ocean and on the Ross continental shelf. The Ross Sea region had just been surveyed by the US research vessel USNS Eltanin during which several thousand kilometres of single channel seismic surveys were taken and used to deﬁne the three main shelf basins (Houtz and Meijer, 1970; Houtz and Davey, 1973). The Eastern Basin was most promising with dipping strata at the basin margin, allowing a drill ship with a subsea ﬂoor penetration limit of around 500 m, to core most of the sedimentary section in three well-chosen sites. 3.3.1. Ship-Based Drilling (1972–1973) The Glomar Challenger sailed south on DSDP Leg 28 from Fremantle, Australia, in December 1972. In the 2 months that followed, it succeeded both to establish the history of continental separation south of Australia and to provide the ﬁrst physical record of Antarctic glaciation extending back to Oligocene times (Hayes, Frakes et al., 1975). The three sites drilled in the eastern Ross Sea (Figs. 3.1 and 3.2) revealed large thicknesses of poorly sorted glacial marine debris ranging in age from late Oligocene (B25 Ma) to Quaternary. The oldest glacial sediments were cored at DSDP Site 270, where they overlay a palaeosol developed on a calc-silicate gneiss basement. DSDP Leg 29, which followed in March 1973, added a new and independent body of data with a signiﬁcant bearing on Antarctic glacial history. The goal of Leg 29 was to sample the ﬂoor of the deep ocean south of Australia and New Zealand for as complete a record as possible of strata representing the whole Cenozoic Era to provide data on ocean history for the new ﬁeld of palaeoceanography. This venture was also successful, both in recovering long and representative stratigraphic sequences and in ﬁnding

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Figure 3.1: Map showing the Antarctic continent and ice drainage systems (based on Drewry, 1983), with sites of the ﬁrst drill holes on the Antarctic continental shelf (DSDP Leg 28). Subsequent shelf sites in the McMurdo Sound region (box – see Fig. 3.3 for detail), Prydz Bay and the Antarctic Peninsula are also shown, each reﬂecting the history of the ice sheet in their respective regions. DSDP Leg 29 drilled lower latitude deep-sea sites south of New Zealand for the ﬁrst Cenozoic isotopic record of ice volume and temperature (see text). Modiﬁed from Barrett (1999), with permission. well-preserved calcareous microfossils for applying a new technique, stable isotope analysis, to these cores (Kennett, Houtz et al., 1975). The result was the ﬁrst record of warm early Cenozoic times followed by a decline in temperature from around early mid-Eocene times (B50 Ma ago,

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Figure 3.2: R/V Glomar Challenger at Site 270 (781S) in January 1973. Seismic section shows how three holes sampled most of the sedimentary section (De Santis et al., 1999).

Shackleton and Kennett, 1975). Two features in the isotope record stand out and are still seen as signiﬁcant today: (i) An abrupt increase in d18O in earliest Oligocene times (now 34 Ma), interpreted as a consequence of global cooling, with extensive sea-ice formation around Antarctica. A major global climate event at this time had already been suspected from the change in terrestrial ﬂora around the world. (ii) Another increase in d18O at B14 Ma, and interpreted to be the result of development of a larger relatively stable Antarctic Ice Sheet like that of today.

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Although deep ocean sediments were useful for their continuous record of past ocean chemistry, and in providing an ice volume–temperature signal from oxygen isotopes, they could not provide information on the extent of ice or regional climate in the high latitudes. This could only come out of sediment cores from the Antarctic margin, where the direct inﬂuence of ice advance and retreat (and perhaps also sea level fall and rise) could be obtained. 3.3.2. Land-Based Drilling (1973–1974) While DSDP was planning its ﬁrst Antarctic and Southern Ocean cruises, the US Ofﬁce of Polar Programs was also decided to address the B180 million years gap in Antarctic geological history, as well as develop a more collaborative approach to major Antarctic science problems. This decision resulted in the launch of the Dry Valley Drilling Project (DVDP), an initiative sponsored by the United States, Japan and New Zealand to explore the late Cenozoic history of the McMurdo Dry Valleys (McGinnis, 1981; Smith, 1981), using a standard ‘‘slim-hole’’ diamond drilling system widely used in mineral exploration. The rig selected was a ‘‘Longyear 44’’, designed to take continuous core of 85, 64 or 48 mm diameter to depths of around 500 m. In the ﬁrst two ﬁeld seasons (1973–1974 and 1974–1975), the project drilled 3 test holes at McMurdo Station, 1 over 300 m deep, and 11 holes in the McMurdo Dry Valleys (Fig. 3.3; Torii, 1981). The most successful of these recovered 320 m of interbedded glacial and interglacial ﬁordal sediments of late Miocene to recent age in the lower Taylor Dry Valley (Powell, 1981; Webb and Wrenn, 1982), indicating several advances and retreats of the ice in that time. Wrenn and Webb (1982) used microfossils in the deposits to provide approximate ages on several geomorphic surfaces in the lower Dry Valleys. 3.3.3. Drilling from Sea Ice (1975) The success of DSDP Leg 28 in the Ross Sea had shown the potential for offshore sediments around Antarctica providing a direct record of the glacial history on land. However, ice ﬂow lines and lithologies of pebbles in cores from the sites in eastern Ross Sea indicated that sediments deposited there were derived from West Antarctica (Fig. 3.1; Barrett, 1975). To ﬁnd a record of the much larger and possibly older East Antarctic Ice Sheet, sites were needed in the western Ross Sea adjacent to the Victoria Land coast.

Figure 3.3: Map and cross-section of McMurdo Sound (adapted from Naish et al., 2008a). (A) The map shows sites for the Dry Valley Drilling Project (1973–1975) and subsequent offshore drilling (MSSTS-1 in 1979, CIROS 2 and 1 in 1984 and 1986, respectively, CRP-1, -2 and -3 in 1997, 1998 and 1999, and ANDRILL MIS and SMS in 2006 and 2007 – details in Table 3.2). Contour interval 200 m. Transantarctic Mountains exposures in light brown, McMurdo Volcanics in dark brown. (B) The cross-section across McMurdo sound shows how a few key drill holes have sampled the entire Cenozoic record from this section of the basin. CRP sites provide an expanded and better dated record 70 km north of the Late Eocene–Early Miocene section cored by CIROS-1.

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A seismic survey in late 1974 showed that the thick sedimentary strata of the Victoria Land Basin did indeed extend close to the East Antarctic coast (Wong and Christoffel, 1981) and provided the justiﬁcation for DVDP15, the ﬁrst Antarctic offshore hole to be drilled from the fast sea ice. Experience in moving equipment 70 km across the sea ice from McMurdo Station on Ross Island to the Marble Point airstrip on the Victoria Land coast had been gained from fuel train traverses, and the same logistics were employed in transporting the DVDP rig and related equipment from McMurdo Station to the site selected for DVDP-15 around 10 km off Marble Point. The rig was set up on 2 m of sea ice in 120 m of water and began drilling in early November 1975 (Fig. 3.4). Logistic delays and drilling problems reduced the time available for drilling, which was terminated at a depth of 62 m with over 52% recovery. The character of the sediments was surprising, for below 12 m of recent glacial debris lay 50 m of basaltic sand, presumed to have been erupted as volcanic ash and then wind-blown across the sea ice before settling (Barrett and Treves, 1981). Nevertheless, the experience had shown the feasibility of drilling from a sea-ice platform.

Figure 3.4: Development of sea-ice drilling systems from 1975 to 1999. (A) DVDP-15, 1975. Water depth 112 m; depth Cored 62 m bsf (below sea ﬂoor). (B) MSSTS-1, 1979. Water depth 195 m; Depth Cored 229 m bsf. (C) CIROS-1, 1986. Water depth 205 m; depth Cored 702 m bsf. (D) CRP-3, 1999. Water depth 305 m; depth Cored 942 m bsf.

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3.4. Developments in Drilling and Thinking in the Late 1970s Although DVDP had concluded its work, the NZ Antarctic Programme was persuaded that a further effort to drill in western McMurdo Sound was justiﬁed. This was named the McMurdo Sound Sediment and Tectonic Studies Project (MSSTS), and led by New Zealand but with US and Japanese scientiﬁc participation. The United States provided the ‘‘Longyear 44’’ rig from DVDP and signiﬁcant logistic support. A camp for the MSSTS-1 hole was set up on the sea ice soon after the late August winter ﬂight to the ice, allowing the drilling to begin in October 1979. Despite the cold and difﬁcult conditions, core was recovered to a depth of 227 m below the sea ﬂoor, when operational problems terminated drilling. After passing through a few tens of metres of Plio-Pleistocene strata, and an interval of virtually no recovery, the strata below 110 m bsf cored well. They provided a record of striking facies changes between diamictites, sand and mud, along with a pollen record and a beech leaf indicating a vegetated coastline and a cold temperate continental margin in those times. Both lithofacies and biofacies were seen as reﬂecting cyclic changes in extent of the continental ice-sheet margin advancing and retreating across the drill site, with associated changes in sea level by tens of metres through glacioeustacy (Barrett and McKelvey, 1986; Barrett et al., 1987). Initially, the strata were thought to extend back to Paleocene times (Webb, 1983) but further study of the sparse faunal and ﬂoral assemblages and recognition of reworked older microfossils led to a late Oligocene age assignment (Webb et al., 1986). While the NZ programme had focussed on drilling in McMurdo Sound, the US programme developed a project to core through the Ross Ice Shelf 420 km from the ocean (Clough and Hansen, 1979). The Ross Ice Shelf Project (RISP) in two successive seasons drilled through 430 m of ice to measure and sample the properties of the 230 m water column, and take cores and photographs of the sea ﬂoor beneath (Webb, 1978, 1979). The sea ﬂoor cores revealed a few tens of centimetres of Late Quaternary mud overlying a metre of midMiocene glaciomarine mudstone with diatomite clasts several millimetres across indicating an interglacial period of ice- and sediment-free biogenic sedimentation at 821S (Webb et al., 1979; Scherer et al., 1988). Terrestrial palynomorphs from the clasts indicate coastal beech forests at this time also. The MSSTS-1 results also provided a glimpse of Antarctic glacial history from the Ross Sea margin, and although there were some issues to be resolved with the drilling technology, scientiﬁc advances in the wider world were supporting the case for a better physical record of past Antarctic climatic events. These included extensive records of seismic stratigraphy from continental margins, along with a new type of analysis that recognized

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sequences developed as a consequence of coastal advance and retreat (Vail et al., 1977). These coastal movements were taken to imply variations in ice volume causing sea level changes of hundreds of metres, and mainly from the mid-Oligocene on. At the same time, the newly developed deep-sea isotope record provided a different basis for inferring ice-sheets on Antarctica, and these indicated an earliest Oligocene initiation (Kennett, 1977). Kennett (1982) incorporated these new concepts and insights from sequence and isotope stratigraphy into a comprehensive synthesis of history and knowledge of the Earth as a system, albeit with a focus on the marine realm, an approach that is now widely accepted and practised. Neither isotopes nor sequence stratigraphy could provide direct evidence of the behaviour and history of the Antarctic Ice Sheet itself, but the ﬁrst coastal Oligocene cores, from MSSTS-1 in 1979, conﬁrmed the cyclic behaviour of ice sheet and sea level (Barrett et al., 1987). Core chronology was not yet adequate to constrain their frequency, but the role of orbital forcing in driving highfrequency Quaternary glacial cycles (Hays et al., 1976) was plainly relevant to early Antarctic Ice Sheets. While signiﬁcant progress was being made in documenting past ice-sheet behaviour, a key event of the 1970s was the publication of John Mercer’s hypothesis for the likely future behaviour of the Antarctic Ice Sheet as a consequence of rising CO2 emissions (Mercer, 1978). He observed that a large area of West Antarctica lay below sea level, the ice sheet being thus inherently unstable, and that much of the ice-sheet margin was buttressed by ice shelves. Observing that present-day ice shelves form only where the January summer isotherm is below 01C, he concluded that a rise of 5–101C, projected as likely in the following 50 years or so, would be sufﬁcient to cause the disintegration of the major ice shelves buttressing the West Antarctic Ice Sheet leading to disintegration of the ice sheet itself. Although there was signiﬁcant disagreement among glaciologists on both the buttressing hypothesis and the immediacy of the threat, the hypothesis provided new impetus for seeking a geological record of the history of both East and West Antarctic Ice Sheets.

3.5. Discoveries Offshore and on the Continent in the 1980s 3.5.1. Further Sea-Ice Drilling – The CIROS Project From 1977 on, the SCAR Working Group on Geology, under the chairmanship of Cam Craddock, provided a useful forum for discussion

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and planning further drilling in McMurdo Sound in an attempt to provide records from the Antarctic margin to resolve the chronology of early icesheet formation and its subsequent evolution. Meetings with its companion group on geophysics were also important for discussions on marine geophysical surveys being planned for the Antarctic margin. The survey by S.P. Lee was particularly useful for future drilling in providing the ﬁrst multichannel section of the western margin of the Victoria Land Basin (Cooper and Davey, 1985; Cooper et al., 1987) and showing a thick seawarddipping sequence that extended close enough to the coast to be drilled from the fast ice. Discussions at the SCAR meeting in Queenstown (Barrett and Webb, 1981) led to a proposal for Cenozoic investigations in the western Ross Sea (CIROS) to drill four holes in two successive seasons from the sea ice off New Harbour and off Granite Harbour (Barrett, 1982), based on the seismic surveys of Wong and Christoffel (1981) and Cooper et al. (1987). The project logistics were managed by the NZ Antarctic Programme with signiﬁcant US support and scientiﬁc participation, and the drill rig and camp were set up in late 1984, 12 km off Marble Point (Fig. 3.3), for the ﬁrst drill hole (CIROS-1). However, concerns grew about the poor state of the sea ice and just prior to drilling, it was decided to shift the camp and rig 20 km inshore to the CIROS-2 site in Ferrar Fiord. After a number of set-backs, CIROS-2 was successfully drilled 165 m below the sea ﬂoor to a granitoid basement, recovering several cycles of Quaternary black sand and diamict and Pliocene mudstone and diamictite. The sequence recorded a number of cycles of glacial advance from the inland ice in the Pliocene followed by cycles in which a Ross Sea ice sheet ﬂowed westward into the Fiord (Barrett and Hambrey, 1992). The core had in fact veriﬁed the ﬁndings from DVDP-11 in 1973, the drilling system had worked and valuable experience had been gained. 3.5.2. Discoveries in the Transantarctic Mountains – The Sirius Formation/Group On land, wet-based glacial deposits at high elevations in the Transantarctic Mountains were providing new evidence of past Antarctic climate. The deposits had been found in the late 1960s, and named the Sirius Formation by Mercer (1972), who saw them as the products of temperate glaciation pre-dating the present ice sheet. However, Mayewski (1975) mapped and analysed more of these deposits, and thought they more likely represented a pre-Pliocene over-riding phase of the East Antarctic Ice Sheet. Later in the

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mid-1980s, yet another view was proposed by Webb et al. (1984) as the result of a laboratory study of samples collected earlier by Mercer, Mayewski and others and reported in detail by Harwood (1986). In this study, samples had been processed from tens of localities along the mountains. Some were barren but others yielded a variety of types and ages of microfossils with some as old as late Cretaceous and a few marine diatoms as young as Pliocene. The latter fossils were thought to provide a maximum age for the Sirius Formation, having been deposited in East Antarctic interior seas and subsequently transported by a growing ice sheet ﬂowed through the mountains at levels more than 1,000 m higher than today’s outlet glaciers. The discovery of plant material, and later leaves and stems, in the deposits (Webb and Harwood, 1987, 1993; Francis and Hill, 1996; Hill et al., 1996) indicated a mean annual temperature almost 201C warmer than today. Mercer’s (1978) case for a collapse of the West Antarctic Ice Sheet through a doubling of atmospheric CO2 had heightened interest in ice-sheet behaviour, but the loss of the East Antarctic Ice Sheet in recent geological times seemed less credible to many, who queried the age of the Sirius Group from recycled diatoms. Indeed, some (e.g. Clapperton and Sugden, 1990) queried the age of the diatoms themselves. However, the Pliocene age of the diatoms was conﬁrmed by dating a volcanic ash in mudstone with the same taxa and cored in CIROS-2 (Barrett et al., 1992).

3.5.3. CIROS-1 – Dynamic Oligocene Ice Sheets Offshore, the most successful drilling for the elusive early Cenozoic preglacial record took place in late 1986. The CIROS-1 drill hole cored to 702 m below the sea ﬂoor with 98% recovery. The core reached back to the late Eocene and spanned the period to the early Miocene, based on magnetostratigraphy underpinned by biostratigraphic datums largely from diatoms, but including foraminifera, palynomorphs and nannofossils (Harwood et al., 1989; Hannah et al., 1997; Wilson et al., 1998; Roberts et al., 2003). However, the record did not quite reach the warm middle Eocene, despite coring the early Oligocene–late Eocene transition in the lower half of the core. The strata were a deep-water turbidite facies with lonestones, indicating some glacial inﬂuence. In contrast, the late Oligocene– early Miocene upper half comprised sequences of diamictite, sand and mud (Hambrey et al., 1989), seen in a much smaller way in MSSTS-1, and typical of near-shore glaciomarine sedimentation with cyclic ice margin advance and retreat and corresponding changes in sea level. The depositional patterns

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clearly showed Antarctic Ice Sheets to be dynamic and reaching the Ross Sea coast from at least late Oligocene times onwards. 3.5.4. Ice-Sheet Proxies from Seismic Stratigraphy and Deep-Sea Isotopes As the results of the MSSTS-1 and CIROS-1 records were being studied, two signiﬁcant papers were published. One was by Haq et al. (1987), a better documented update of the Vail et al. (1977) monograph with a more detailed sea level curve based on coastal advance and retreat histories extracted from seismic imaging of continental margins around the world. It still showed a sharp and substantial mid-Oligocene shallowing, ostensibly from Antarctic Ice Sheet formation, along with some isolated shoaling events in the Paleogene and Cretaceous. The paper emphasized the cyclic sea level changes with a frequency of 1–3 million years, third-order cycles, inferred to be glacioeustatic. This link between ice sheets and sea level was invoked to explain cyclic patterns in Cenozoic strata in Australasia by Loutit and Kennett (1981). The other paper was a compilation of deep-sea isotope data from numerous DSDP and Ocean Drilling Program (ODP) sites in the Paciﬁc and Atlantic Oceans by Miller et al. (1987) – the smoothed curves for both oceans were a little different in detail but their main features included both the earliest Oligocene and middle Miocene positive isotope shifts ﬁrst seen by Shackleton and Kennett (1975). At the time, no cyclicity comparable to that proposed by Haq et al. (1987) was recognized in the deep-sea isotope record, but the record plainly indicated continental ice build-up in the earliest Oligocene, not the mid-Oligocene indicated by the Haq et al. (1987) analysis. A comparison of the two curves is shown in Fig. 3.5. Which was correct? CIROS-1 cores had the oldest signiﬁcant diamictites in the late Oligocene, but the Transantarctic Mountains, with their origins around 55 Ma (Gleadow and Fitzgerald, 1987), could well have been a barrier to earlier continental ice. 3.5.5. DSDP–ODP Drilling Dates Initiation of Antarctic Ice Sheet By this time, the ODP had approved two further Antarctic legs, one in the Weddell Sea (DSDP 113) and one in Prydz Bay (ODP 119), seeking a glacial history from other sectors of the Antarctic margin. The Weddell Sea leg cored thick continuous Paleogene deep-sea sediments between 601 and 711S, recording a relatively warm climate, cooling sufﬁciently for cold deep water formation and some limited ice to form on land in the early Oligocene (Barker, Kennett et al., 1988). However, ODP 119 was able to drill several

A History of Antarctic Cenozoic Glaciation

51

Figure 3.5: A comparison of the deep-sea isotope record of Miller et al. (1987), the eustatic sea level curve of Haq et al. (1987), and the record from cores from the Eastern and Victoria Land Basins and Prydz Bay (from Barrett and Davey, 1992). Both the isotope record and the eustatic sea level curve were considered by their proponents to be sound Ice volume proxies, but showed signiﬁcant differences in early Cenozoic times. Reproduced from Barrett and Davey (1992), with permission from the Royal Society of New Zealand.

52

P. Barrett

shelf sites in Prydz Bay, coring offshore-dipping strata, ranging from Cretaceous terrestrial sediments near the coast to thick late Eocene–early Oligocene diamictite mid-shelf to Neogene diamictites in the outer shelf (Hambrey et al., 1991). No better age has been extracted from these fossilpoor sediments, but the result was sufﬁcient to make it clear that continentalscale ice sheets on Antarctica were delivering ice to the shelf edge beyond the limits of the modern ice sheet at least from early Oligocene times. ODP Leg 120 that followed was able to both conﬁrm this physically through the occurrence of ice-rafted debris at the Eocene–Oligocene boundary and provide a better deep-sea isotope record from continuous Paleogene cores from the edge of the Kerguelen Plateau to the north (Zachos et al., 1992). The timing of the ﬁrst of the big Antarctic-wide ice sheets had now been established (Wise et al., 1991). 3.5.6. Planning Future Research through SCAR The 1980s also saw work on Antarctic glacial history further facilitated by SCAR, initially through a Workshop convened by Peter Webb at the SCAR meeting in San Diego (CA) in 1986 on Cenozoic Palaeoenvironments in Southern High Latitudes. From that meeting emerged a plan for a Group of Specialists to coordinate and promote research on this broad topic. This was approved by SCAR and the Group began work with a workshop on polar drilling in Columbus, OH, in 1988. This meeting reviewed a range of drilling technologies from shelf ice, sea ice and ships, as well as shallow ship-based coring. Extensive ship-based coring programmes from the 1960s using the USNS Eltanin, USCGC Glacier and more recently the R/V N.B. Palmer provided some hard won data, but typically could not penetrate beyond the hard diamictites of the Last Glacial Maximum (with rare notable exceptions, such as the Cretaceous strata cored in Ross Sea by Domack et al., 1980). Webb (1990) provides a useful review of the state of knowledge of Antarctic glacial history at this time. At the same 1986 San Diego meeting, there was also interest in the growing body of seismic data being collected around the Antarctic margin, because of both its scientiﬁc value for Antarctic glacial history and also the political sensitivity to possible misuse. This led to the formation of the ancillary Antarctic Offshore Stratigraphy (ANTOSTRAT) Project, led by Alan Cooper (see Chapter 5). By 1990, the project had ﬁve regional working groups to organize seismic data from the ﬁve regions of the Antarctic margin, and make it available through a Seismic Data Library System (Cooper and Brancolini, 1997). However, the main goal was to use these data

A History of Antarctic Cenozoic Glaciation

53

as the basis for coring proposals in key areas in each region in the next decade under the aegis of the ODP. 3.5.7. Insights from Deep Ice Cores While the Antarctic drilling community was now focussed on a record of icesheet history going back to the Eocene times, the ice coring community was providing a remarkable climate record of the last glacial cycle with measurements of the atmospheric gas composition and proxy temperature history from the Vostok core (Barnola et al., 1987). The saw-tooth pattern of the glacial cycle, developing slowly (tens of thousands of years) with irregularities from varied orbital forcing, and deglaciation taking place relatively quickly (thousands of years), showed a close association with eustatic sea level change for the same period (Fairbanks, 1989). This gave new impetus to the importance of understanding Antarctic Cenozoic glacial history in order to appreciate its likely response to rising CO2 emissions, which by then was coming to be seen by senior climate scientists as an imminent danger (e.g. Hansen et al., 1988; Schneider, 1989).

3.6. Advances in the 1990s The 1990s saw the publication of several major reviews on the state of knowledge of Antarctic climate history, notably symposium volumes by Kennett and Warnke (1992, 1993), ANTOSTRAT monographs on the Ross Sea (Cooper, this volume) and in support of ODP drilling, three chapters in Tingey’s (1991) monograph on the geology of Antarctica (Anderson, 1991; Denton et al., 1991; McKelvey, 1991) and the review of sedimentation on the Antarctic continental shelf by Hambrey et al. (1992). By now, the broad chronological framework had been established through drilling in McMurdo Sound (CIROS-1, Barrett, 1989) and in Prydz Bay (Barron, Larsen et al., 1989) on opposite sides of the continent. However, neither the earliest Oligocene onset of glaciation nor the middle Miocene transition had been sampled by drilling on the continental shelf, and progress was hampered by a lack of chronological tools comparable with those used to date deep-sea sediments (e.g. microfossil datums based on abundant rapidly evolving taxa integrated with magnetic reversal stratigraphy in continuous sedimentary sequences). ANTOSTRAT contributed to a Detailed Planning Group of the ODP at College Station (TX) in 1994 for planning proposals for all ﬁve ANTOSTRAT sectors. However, in the event, only two legs were approved

54

P. Barrett

for drilling (Legs 178 in 1998 and Leg 188 in 2000). Both had some success in ﬁlling out the glacial chronology in each region in cores from fans and drifts on the slope and rise (Barker, Camerlenghi et al., 1999; O’Brien, Cooper et al., 2001; Cooper, this volume), though poor recovery marred the success of drilling on the continental shelf. 3.6.1. Antarctic Sea-Ice Drilling Progress in further sea-ice drilling began slowly with a SCAR workshop in Bremerhaven in 1990 for discussions that led to the formation of SCAR Group of Specialists on Global Change (GLOCHANT). While Cenozoic glacial history was seen as beyond the remit of the group, the meeting was productive in that it began a discussion among US, NZ, Italian, German and UK scientists for a workshop in Wellington (Barrett and Davey, 1992) that initiated the Cape Roberts Project (CRP). Following the meeting, national programmes agreed to plan for drilling four holes in two seasons to core a 1,500 m sequence that was thought from seismic correlation to extend the CIROS-1 record back to the Cretaceous. A new drilling system was also needed (Fig. 3.6).

Figure 3.6: Diagram showing the advances made in sea riser design after the ﬁrst decade of McMurdo Sound drilling for the Cape Roberts Project, and the further developments that were required for ice-shelf drilling (Modiﬁed from Harwood et al., 2003, with permission from the ANDRILL Science Management Ofﬁce).

A History of Antarctic Cenozoic Glaciation

55

The Cape Roberts Project sites were drilled, after a year’s delay on account of poor ice conditions, and with three drilling seasons (1997–1998 to 1999–2000, reported in CRST (Cape Roberts Science Team) (1998, 1999, 2000) and Davey et al. (2001). The 1,500 m-thick section cored provided the ﬁrst proximal high-resolution record from the Antarctic continental shelf of climate history for the period from 34 to 17 Ma. A combination of biomagnetostratigraphic dating with Ar–Ar radiometric ages at several key points provided robust (o0.5 million years) resolution for most of the interval and excellent resolution (o0.1 million years) for three cycles around 24 Ma ago (latest Oligocene) (Florindo et al., 2005, but see update by Naish et al., 2008b). These showed for the ﬁrst time that the Antarctic Ice Sheet was responding to orbital forcing with 40,000 and 100,000 year frequencies in the distant past (Naish et al., 2001). The cores also recorded the inﬂuence of orbital forcing on Antarctic Ice Sheets through Oligocene and early Miocene times as well as persistent slight cooling over this period (Barrett, 2007; Dunbar et al., 2008). This contrasted with the late Oligocene warming deduced from a review of deep-sea oxygen isotope data by Zachos et al. (2001), but was consistent with the reinterpretation of oxygen isotope data sets for this period by Pekar et al. (2006). 3.6.2. Ice Sheet and Climate Modelling ANTOSTRAT in a few short years had greatly extended the signiﬁcance and value of the few drill holes that could be drilled on the Antarctic margin by providing a basis for correlating events from basin to basin. However, the behaviour of the ice that eroded and deposited the sediments was not well understood, especially on a continental scale. It was assumed that on a roughly circular continent, the extent of the ice would be similar in all directions, but Antarctica has enough irregularities, mountains and basins to see that local geological histories could vary. Ice-sheet modelling provided a means of objectively experimenting on and visualizing ice-sheet behaviour, Oerlemans (1982) providing an early example. Subsequent ice-sheet modelling by Huybrechts (1993) (Fig. 3.7) was signiﬁcant because it showed a clear and consistent relationship between temperature and ice volume, including a slight increase in ice mass with initial warming, thus linking past Antarctic temperatures and sea level change. It indicated the loss of West Antarctic ice with a regional warming of 101C, in the range Mercer (1978) had suggested, and provided patterns of ice growth and decay that varied around the continent with changes in temperature. The model also supported the view that the earliest ice sheets

56

P. Barrett

Figure 3.7: Modelling the Antarctic Ice Sheet for progressively higher regional temperatures (Huybrechts, 1993). If a polar ampliﬁcation factor of 2 is assumed (Manabe and Stouffer, 1980), then the effective rise in average global temperature to achieve each of these states would be 1/2 of the temperature rise shown.

would have reached the coast in the Prydz Bay, whereas the Transantarctic Mountains offered a signiﬁcant barrier ice reaching to the Ross Sea. 3.6.3. Discoveries in the Transantarctic Mountains Reinterpreted Huybrechts’ model also suggested that it was not possible to create conditions in which an ice sheet could grow and over-top the present Transantarctic Mountains at their present height, depositing wet-based

A History of Antarctic Cenozoic Glaciation

57

glacial debris at the margins. Some argued that the mountains were lower in the Pliocene and had risen rapidly since then (Behrendt and Cooper, 1991), but this was subsequently countered in a review of Antarctic tectonic history by Fitzgerald (2002). However, there were other substantive concerns. New high-resolution deep-sea isotope records of the last few million years gave no indication of negative anomalies to be expected from extensive ice-sheet recession implied by Pliocene marine basins in the East Antarctic interior (Denton et al., 1991; Kennett and Hodell, 1993). Marchant et al. (1993) reported volcanic ashes at high elevations in the McMurdo Dry Valleys and dating back to B11 million years, whose survival seemed inconsistent with a warm Pliocene and an ice-free Antarctic interior. Also, diatoms were discovered in South Pole snow (Kellogg and Kellogg, 1996) and surﬁcial rock debris, and counts showed they were extremely rare in Sirius deposits (Barrett et al., 1997). These observations led to the view that the Sirius deposits predated the widely accepted mid-Miocene cooling, the few agecritical Pliocene diatoms resulting from atmospheric transport. The report by Gersonde et al. (1997) of a Pliocene meteorite impact event ejecting diatomaceous sediment into the stratosphere from the southeastern Paciﬁc Ocean ﬂoor provided another source for possible atmospheric contamination. However, Harwood and Webb (1998) maintained that the marine diatoms and diatomite clasts found by Harwood were larger than winds could carry. In reviewing Antarctic climate through Cenozoic times, Barrett (1996, 1999) noted that McMurdo Sound drill cores had shown Oligocene and early Miocene Antarctic Ice Sheets were of similar extent to those of recent times during glacial periods, with largely ice-free forested coasts during interglacial periods, thus resembling the dynamic behaviour of Northern Hemisphere ice sheets in the Quaternary. This contrasted with the relative stability of the Antarctic Ice Sheet at this time, when there has been little change in its central elevation, although a signiﬁcant volume of ice has been lost around the margin, B15 m of sea level equivalent (SLE) since the Late Glacial Maximum (Zwartz et al., 1997). While offshore mid-Miocene to early Pliocene strata had yet to be cored, seismic stratigraphy of the continental shelf showed numerous sediment packages and unconformities indicating signiﬁcant glacial deposition and erosion extending to the shelf edge. These data implied signiﬁcant dynamism of the ice sheet at the margins (Anderson and Bartek, 1992), but indicated little about its interior health. The evidence cited above for a persistent late Neogene ice sheet in the East Antarctic interior was not necessarily inconsistent with clear evidence of Pliocene warmth from marginal coastal marine sediments, as in the region east of Prydz Bay (Harwood et al., 2000).

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3.7. Advances in the First Decade of the Twenty-First Century By the mid-1990s with the work of the ANTOSTRAT regional groups coming to a close, and planning well advanced for further geological data gathering by ODP and sea-ice drilling (Cape Roberts Project), the value of modelling as a means of illustrating geologically based scenarios, as well as using geological information to constrain models, seemed increasingly obvious. This was the theme of ANTOSTRAT workshops in Hobart (AU) in 1997 (Cooper and Webb, 1998; Webb and Cooper, 1998), Wellington (NZ) in 1999 and Erice (IT) in 2001 (Cooper et al., 2002), with the latter providing the starting point for the current SCAR Scientiﬁc Programme on Antarctic Climate Evolution (ACE).

3.7.1. Ice–Atmosphere–Ocean Modelling Modelling the effect of atmospheric CO2 on the sensitivity of the Antarctic Ice Sheet to orbital forcing by DeConto and Pollard (2003), and how this may have initiated Antarctic Ice Sheet formation, was a turning point in understanding past ice-sheet behaviour. For three decades, it was believed that the main cooling inﬂuence came from thermal isolation with the formation of oceanic ‘‘gateways’’ (Kennett, 1977). The modelling showed, albeit simply and with signiﬁcant assumptions, that B80% of the inﬂuence could have been due to the cooling of atmospheric temperature from declining atmospheric CO2 levels. This has gained support from evidence of a 1 km drop in the carbonate compensation depth in the latest Eocene (Coxall et al., 2005) and a proxy record based on d13C from alkenones indicating an irregular decline in atmospheric CO2 levels from more than 1,000 ppm to less than 400 ppm between 40 and 24 Ma (Pagani et al., 2005). Changes in ocean circulation through opening gateways have also been accommodated (Huber and Nof, 2006). As a result of this group of studies, the signiﬁcance of CO2 as a greenhouse gas inﬂuencing Antarctic climate, ice sheets and ocean chemistry has been further enhanced.

3.7.2. Landscape Modelling A different application of modelling has developed in the last few years to explore the response of a particular sector of the ice sheet to changes on geological time scales. Taylor et al. (2004) studied the Lambert drainage

A History of Antarctic Cenozoic Glaciation

59

basin on account of its signiﬁcant size (16% of the area of the Antarctic Ice Sheet) and the fair degree of knowledge of its both bedrock topography and geological history. They found that once the climate had cooled to the point of initiating glaciation, the erosional history of the Lambert Basin became strongly inﬂuenced by continuing graben development, with past erosive deepening limiting the future extent of the ice margin regardless of the climatic regime. Jamieson et al. (2005) used a similar but broader analytical approach to reconstruct the Lambert drainage basin landscape since the ﬁrst ice sheets developed 34 Ma ago, and found the relative inﬂuence of ﬂuvial and glacial erosion through this time to be similar, with only slight glacial modiﬁcation to long established ﬂuvial drainage patterns. Jamieson and Sugden (2008) extended this approach to the landscape evolution of the whole continent with a similar result. They acknowledge the paucity of constraining geological evidence, and the lack of account taken of changes in area or elevation of various sectors on account of tectonism. However, the potential for development is plain, and a link with climate-oriented modelling will be of crucial importance.

3.7.3. Sea- and Shelf-Ice Drilling Extracting a direct geological record of climate and ice-sheet extent from offshore drilling in McMurdo Sound has continued on the back of the Cape Roberts Project. The successor project, ANDRILL, was designed to sample an expanded Neogene section with a pair of sites, one for the latest Neogene and the other for the middle Miocene transition. There was to be an added twist – the ﬁrst site would be in the centre of the Victoria Land Basin and located on the ice shelf just south of Ross Island to obtain a history of the Ross Ice Shelf/expanded West Antarctic Ice Sheet cycle through late Neogene times. This required further technological development for maintaining an access hole through an ice shelf in deeper water and with deeper coring sought (Fig. 3.6). Both holes have now been successfully completed with excellent recovery (Naish et al., 2008a; Harwood et al., in press; Table 3.2). They also overlap with each other and CRP-1 (Fig. 3.8), and provide record of numerous cyclic ice-sheet ﬂuctuations under varying climatic regimes over the last B18 million years. Perhaps the most interesting and unexpected result has been recovery of the 90 m of Pliocene diatomite in AND-1B, indicating open water in the Ross Embayment (and most likely throughout West Antarctica) for several hundred thousand years when global temperatures were only 2–31C

1984

1997 1998 1999

CIROS

CRP

1 2 3

2

1 1

77100uS 77100uS 77100uS

77141uS

77134uS 77105uS

77126uS

McMurdo Sound area – offshore DVDP 15 1975 15

1979 1986

77150uS 77151uS 77151uS 77135uS 77135uS 77138uS

McMurdo Sound area – onshore DVDP 1973 1 1973 2 1973 3 1974 10 1974 11 1974 12

MSSTS CIROS

77126uS 77126uS 77126uS

271 272 273

Latitude

77126uS

1973

Ross Sea DSDP 28

Site

270

Year

Project

163145uE 163143uE 163143uE

163132uE

163123uE 164130uE

164123uE

166140uE 166140uE 166140uE 163131uE 163125uE 162151uE

178130uW 178130uW 178130uW

178130uW

Longitude

154 178 295

211

195 197

122

67 47 48 3 80.2 75.1

562 629 491

634

Elevation/ depth (7) (m)

148 624 939

168

230 702

62

201 179 381 182 328 185

233 439 333

423

Depth cored (m)

86 95 97

67

62 98

52

98 96 90 83 94 98

7 37 25

62

Recovery (%)

Diamict – early Miocene Mudstone – Oligocene Sandstone – Devonian

Mudstone – late Oligocene Boulder conglomerate – late Eocene Gneiss – early Paleozoic

Black sand – early Pleistocene

Basalt – Late Quaternary Basalt – Late Quaternary Basalt – Late Quaternary Diamict – late Miocene Diamict – late Miocene Migmatite – early Paleozoic

Diamict clasts – early Pliocene Diamict – early Miocene Diamict – early Miocene

Gneiss – early Paleozoic

Oldest core

Table 3.2: Antarctic coastal and continental shelf rock-drilling sites, 1973 to 2007.

Barrett and Hambrey (1992) CRST (1998) CRST (1999) CRST (2000)

Barrett and Treves (1981) Barrett (1986) Barrett (1989)

Kyle (1981) Kyle (1981) Kyle (1981) Powell (1981) Powell (1981) Powell (1981)

Hayes, Frakes et al. (1975)

References

60 P. Barrett

63151uS

63115uS 63120uS 63116uS

3

5 6 12

Note: ?-indicates age given is uncertain.

2006

64152uS 64157uS 66153uS 66148uS 64100uS 62117uS

1098 1099 1100 1102 1103 1

SHALDRIL

66124uS

1097

Antarctic Peninsula ODP 178 1998

2005

66155uS 67142uS

743 1166

2000

SHALDRIL

67133S

742

ODP 188

68141uS 68123uS

740 741

77146uS

77155uS

67117uS

2

1

739

1988

2007

ANDRILL

Prydz Bay ODP 119

2006

ANDRILL

52122uW 52122uW 52150uW

54139uW

64112uW 64119uW 65142uW 65151uW 65128uW 58145uW

70145uW

74142uE 74147uE

75124uE

76143uE 76123uE

75105uE

165117uE

167101uE

506 532 442

340

1010 1400 459 431 494 488

563

989 475

416

808 551

412

380

840

23 21 4

20

47 108 111 15 363 108

437

97 381

316

226 128

487

1139

1285

40 6 64

32

99 102 5 6 12 87

14

22 19

53

32 26

34

98

98

Naish et al. (2008a) Harwood et al. (in press)

Mudstone – late Eocene/early Oligocene Muddy sand – mid-Miocene Muddy sand – early Pliocene Mudstone – Oligocene

Mud – Holocene Mud – Holocene Diamict – Pleistocene Diamict – Pleistocene? Diamict – late Miocene Mud – Late Pleistocene

Diamict – early Pliocene

http://shaldril. rice.edu/ Anderson et al. (2007)

Barker, Camerlenghi et al. (1999)

Diamict – late Eocene–early Barron, Larsen Oligocene et al. (1989) Red beds – ?Triassic Sandstone, siltstone – ?early Cretaceous Mudstone, diamict – ?Eocene– Oligocene Diamict – Pleistocene Claystone – late Cretaceous O’Brien, Cooper et al. (2001)

Diamict – late Oligocene

Basalt – early Miocene

A History of Antarctic Cenozoic Glaciation 61

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P. Barrett

Figure 3.8: Cenozoic time–temperature graph showing the time intervals now covered in moderate detail by McMurdo Sound Drilling. ODP Legs 119 and 188 covered the same interval but with signiﬁcant time gaps and lower core recovery. Intervals covered by much higher resolution deep ice cores and the continuous record from multiple sites in deep-sea sediment cores are shown for comparison. The temperature curve from Crowley and Kim (1995), which is based on the isotope summary of Miller et al. (1987), is modiﬁed to show the effect of the methane discharge at 55 Ma (Zachos et al., 2003).

higher and CO2 levels little more than those of today (Van Der Burgh et al., 1993; Ravelo et al., 2004). The middle Miocene climatic optimimum has also been cored at the Antarctic margin for the ﬁrst time (AND-2, Harwood et al., in press). Much remains to be done for a detailed analysis of sedimentary features, fossil assemblages, physical, chemical and magnetic properties and geochronology for comparing this record and that of AND-1B with those of deep-sea and coastal records in lower latitudes. There is also the prospect for the ﬁrst time of correlating the last million years of the AND-1B core with the

A History of Antarctic Cenozoic Glaciation

63

temperature and atmospheric CO2 record of the last eight glacial cycles in the ice core from Dome C (EPICA, 2004).

3.7.4. New Discoveries in the Transantarctic Mountains Another remarkable record of the mid-Miocene climate transition has recently been reported from on land in the Transantarctic Mountains west of the McMurdo Dry Valleys (Lewis et al., 2007). Polar conditions since the midMiocene have preserved wet-based ice marginal sediments overlain by thin dry-based glacial deposits between 1,200 and 1,500 m above sea level. Both lithofacies include volcanic ash layers dating the transition at between 13.94 and 13.62 Ma. A different but related study involved dating similar volcanic ash layers in sediments interpreted as ﬂood deposits in the Labyrinth, a set of spectacular anastomosing channels formed within a dolerite sill in the ﬂoor of upper Wright Valley nearby (Lewis et al., 2006). The ages constrain the hypothesized megaﬂood event(s) to within the period 12.4 and 14.4 Ma, and also support the concept developed by Denton and Sugden (2005) that megaﬂoods may have had a signiﬁcant role in the past behaviour and erosional effectiveness of the Antarctic Ice Sheet in earlier times. Miocene megadebris ﬂow deposits up to 400 m thick and over 200 km in extent on the slope and rise off Wilkes Land (Donda et al., 2007) might also be related. Recent reports (Wingham et al., 2006) suggest that signiﬁcant subglacial ﬂows are still taking place beneath the Antarctic Ice Sheet, though perhaps not on that scale.

3.7.5. Subglacial Lake Sediments Sought Finally, advances in the search for evidence relating to ACE in this decade have included glaciological and geophysical surveys investigating the many lakes beneath the Antarctic Ice Sheet (Siegert et al., 2005) with a view to using sediments deposited in them as recorders of the basal history of the ice sheet over the period of their existence (Mayer et al., 2007).

3.8. Future Prospects for Improving Knowledge of the History of the Antarctic Ice Sheet The activities described above indicate continuing gains in signiﬁcant new knowledge and a more broadly consistent understanding of Antarctic

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ice-sheet history through improved dating resolution, more careful ﬁeld observations on land and more advanced geochemical techniques and drilling technology offshore. Understanding present-day ice-sheet behaviour is rapidly advancing through modelling studies integrated with ground truth for surface velocities and mass balance from remote sensing, ice core climatology data from ITASE (2007) and ongoing climate monitoring (AGCS, 2008). However, high-resolution data on geological time scales are still needed from key sectors of the Antarctic Ice Sheet to understand its past behaviour as a basis for credible projections for the future. 3.8.1. Future Sea- and Shelf-Ice Drilling The history of drilling on the Antarctic continental shelf since 1973 is summarized in Fig. 3.9. Both ship-based and ﬂoating ice-based operations typically take 5–15 years from conception to execution and require the involvement of a committed group of scientists to execute successfully. The excellent core quality and recovery (95–98%) has come about because of the use of minerals industry technology, which is designed for optimal core recovery, and a sea riser that allows the circulation of a ﬂuid of the right density and viscosity to remove cuttings and protect hole and core. This technique has now been used to drill over 1,200 m beneath the sea ﬂoor and in water almost 1,000 m deep. The ability to re-enter a pre-existing hole, which would be potentially useful for drilling beneath rapidly moving ice shelves, has yet to be developed, but considered feasible (A. Pyne, personal communication, 2008). This may well be useful for ANDRILL’s proposed site on the ice shelf east of Ross Island to sample through early Cenozoic strata and capture the transition from the ‘‘greenhouse’’ world on the Antarctic margin (Decesari et al., 2004). 3.8.2. Future Ship-Based Drilling Core recovery from ship-based drilling on the Antarctic continental shelf is typically poor (a few to 30%) in unconsolidated diamicton, but can be high (W95%) in calm seas and either ﬁne-grained mud lacking lonestones or lithiﬁed sediment (including diamictite). Sites in deeper water (W400 m) are easier for operational reasons, but depth penetration has rarely been much more than 400 m in practice, on account of drill bit failure. However, hole re-entry and drill bit replacement is possible. Ship-based drilling is normally subject to more ice and weather constraints than ﬂoating sea-ice or ice-shelf

A History of Antarctic Cenozoic Glaciation

65

Figure 3.9: Cumulative metres of sedimentary strata drilled on the Antarctic continental shelf since drilling began in 1973, based on data from Table 3.2. Note that only the ODP holes drilled on the shelf are shown here – signiﬁcant holes were drilled on the slope and rise also. The graph shows the episodic nature of progress in recovering core, reﬂecting the time and effort required from a relatively small community in both developing proposals and projects and then seeing them through to publication. The better recovery from ‘‘Slimhole’’ drilling is a consequence of the ability to recirculate drilling ﬂuid, a smaller hole diameter and a stable drilling platform.

drilling. The upcoming Integrated Ocean Drilling Program (IODP) Expedition 323, which plans to occupy both a Neogene and a Holocene site on the continental shelf off Wilkes Land in early 2009, will be a useful test of technology and experience gained from previous legs (IODP, 2007). It will also provide the ﬁrst long time scale geological history for this sector of the Antarctic margin, of interest because it will provide a test for the histories developed separately from drilling in the McMurdo region of the Ross Sea and in Prydz Bay. Despite the quality and length of core it yields, the ﬂoating ice technology can be used only in coastal or inland locations, leaving open huge areas of the

66

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shelf that are now extensively covered with seismic surveys. Considerable effort in the past three decades has gone into ship-based coring programmes to penetrate the layer of compact diamictite deposited when the continental shelf was largely covered with grounded ice during the Last Glacial Maximum, but with little success. SHALDRIL, a US Antarctic initiative (http://www.arf.fsu. edu/projectsShaldril.cfm), has been a recent development to address this need and help correlate the extensive seismic stratigraphy between the various sectors of the Antarctic continental shelf with other regions and other continental margins. The record thus far (Table 3.2) indicates that further technological development is needed for successful sampling.

3.8.3. Time Frame and Context for Future Ice-Sheet Investigations The need for soundly based projections for the future behaviour of the Antarctic Ice Sheet is now critical. While ice-sheet modelling and remotesensing measurements from satellites and aircraft are likely to lead this ﬁeld, palaeoclimate records from ancient ice and sediment are also crucial for documenting ice-sheet response and constraining models for a warmer world. The ANDRILL Pliocene palaeoclimate record with its indications of much reduced ice on West Antarctica on many occasions alerts us to the risks of a world in which atmospheric CO2 levels are o500 ppm (Royer, 2006). However, more high-quality palaeorecords are needed close to all major sectors of the ice sheet. Future sites might be considered not only for their utility in adding to the jigsaw of Antarctic glacial history, but also how the results might be used to test or constrain future modelling and to inform the IPCC assessment process in its effort to integrate results from the whole Earth climate system. The last year has seen increased concern over the large uncertainties in estimates of the loss of Antarctic ice mass from climate change in the last decade, and the issue is complex with estimates made by different methods and over different time frames (review in AGCS, 2008). The broad pattern is of net accumulation over the dome of central East Antarctica with loss of ice around the Antarctic margin, especially the Paciﬁc Coast of West Antarctica, though how much results from long-term trends compared with recent rises in temperature is unclear. An estimate made for Meehl et al. (2007) was around 0.270.35 mm/year of sea level equivalent (SLE) (Lemke et al., 2007), but a more recent report estimates an annual mass loss by 2006 equivalent to 0.470.2 mm SLE (Rignot et al., 2008), almost twice the value estimated for 1996.

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Figure 3.10: History of Antarctic Climate Evolution in the context of changes in global average temperature from 1900 to the present and projected to 2100 (Meehl et al., 2007). The coloured lines beyond 2000 represent possible future global temperature paths from different greenhouse gas emission scenarios. Dashed lines show ‘‘pre-industrial’’ and þ21C temperature levels (European Commission, 2007). Drilling key Antarctic locations for high-quality geological records of past ice-sheet behaviour will help constrain models and reduce uncertainties. Research results will need to be published within the next 4 years to contribute to IPCC AR5 and within the next decade (red dashed box) to contribute to IPCC AR6. Figure 3.10 sets the progress in Antarctic geological drilling against the IPCC review process and the projected rise in global average temperature for various future scenarios. The diagram suggests that only one or two major ANDRILL-type projects in the next decade will yield results in time to have a signiﬁcant inﬂuence on world’s climate community and the public at large while there is still a window of opportunity to mitigate the worst effects of climate change on polar ice sheets (Hansen et al., 2005, 2008). Parallel planning and increased resourcing for site surveys will be especially important for prospective sites beneath ice shelves, which are virtually unexplored and techniques are new, slow and laborious. Glaciological research over the last decade has shown that the Antarctic Ice Sheet can be expected to respond quite differently in different sectors and on different time frames (Vaughan, 2005). Areas of potentially signiﬁcant ice

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Figure 3.11: Velocity map of Antarctic Ice Sheet showing the pattern of present-day ice ﬂow and the main ice drainage systems. The areas for which ice-sheet history are best known – McMurdo Sound (MS) and Prydz Bay (PB) – are circled in brown. Drilling off Wilkes Land (WL) in the dashed brown circle should provide a parallel history from this sector in 2009. The red circles in Pine Island Bay (PIB), the Filchner–Ronne Ice Shelf (FRIS), the Siple Coast (SC) and the Totten Glacier (TG) represent some areas of potential interest for future drilling to seek records of ice-sheet response at the margin to climate change in warmer times in the geological past. Reproduced from Bamber et al. (2000), with permission. loss and whose past history are poorly known are shown in Fig. 3.11. The Paciﬁc coast of West Antarctica, especially Pine Island Bay, has long been recognized as vulnerable to ice loss (Hughes, 1981). Reports of active subglacial hydrological systems beneath the Recovery Glacier, which feeds the Filchner Ice Shelf (Bell et al., 2007), the Totten Glacier in Wilkes Land (Rignot and Jacobs, 2002) and West Antarctica’s Siple Coast (Fricker et al., 2007) indicate the importance of assessing ice-sheet response to climate change in these sectors also. These are also areas of relatively rapid ice ﬂow, though not beyond the ability of a drill system on ﬂoating ice to handle. High-resolution palaeoclimate records dating back at least through Pliocene times in each sector would be helpful. Wilkes Land is already covered with the scheduled IODP Wilkes Leg. Potential drill sites in Pine Island Bay, on the Filchner Ice Shelf and on the Siple coast would all provide challenges

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both for site surveys, logistics and drilling, but the results could well more than justify the effort and cost. The selection of targets for geological drilling for Antarctic climate history, and the assembly of science teams to carry out the work, has until recently been relatively straightforward – so little was known and the pressures derived from our curiosity. Today, the wider community is expecting a much greater understanding of the Antarctic ice–atmosphere– ocean system, so that the threat to future ice-sheet stability can be reliably assessed. One of the key tasks will be identifying, planning, surveying and coring a few critical sites for high-quality records of past ice-sheet behaviour in the most vulnerable sectors of the Antarctic as a guide to modelling icesheet response to future greenhouse gas levels.

ACKNOWLEDGEMENTS I wish to thank a number of people for their help and companionship through the last decades of this history. First and foremost is Peter Webb, who with Barrie McKelvey was the ﬁrst Antarctic expedition from Victoria University of Wellington. I ﬁrst met Peter in 1968, when he visited Colin Bull, leader of the second VUW Antarctic expedition, but then Director of the Institute of Polar Studies, Ohio State University, when I was a graduate student there in 1968. Peter and I subsequently spent 10 weeks in late 1972 and early 1973 together on DSDP Leg 28, drilling in the Southern Ocean and Ross Sea, and from that experience and Peter’s involvement in the Dry Valley Drilling Project, the concept of drilling for Antarctic glacial history from the sea ice was born. Peter led the SCAR Cenozoic Group of Specialists (1986–1996) and the US component of McMurdo Sound drilling, and his contribution to Antarctic glacial history continues to this day. I am also grateful to Alex Pyne for his long-standing appreciation of the scientiﬁc goals of Ross Sea drilling and his technical skills, insight and commitment to the development of drilling technology and logistic support. This began with his observations as a graduate student and drill site science manager for the MSSTS-1 drill hole in 1979 and continued through the CIROS and Cape Roberts Projects to his development of the remarkable ANDRILL system, including its hot water drilling component for ice shelf operations. Key ﬁgures from the early days include Bob Clark, Professor of Geology at VUW from 1954 to 1985, who in 1957 initiated the University’s annual expeditions and kept them going until my arrival there in 1970, Sam Treves, University of Nebraska, lead PI for DVDP-15 in 1975, the ﬁrst attempt at

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offshore drilling from the sea ice, and Trevor Hatherton and Bob Thomson, who as head of Geophysics Division and Antarctic Division, DSIR, respectively, provided practical support and approval for the next three attempts at sea-ice drilling, ﬁnally leading to the success of CIROS-1 in 1986. From the late 1980s on, Alan Cooper and the ANTOSTRAT group broadened my awareness of marine geophysics, and about the same time my appointment to the SCAR Group of Specialists on Environmental Affairs and Conservation, led initially by Nigel Bonner and then David Walton, was signiﬁcant in helping me deal with the environmental implications of offshore drilling from both practical and scientiﬁc perspectives. In the 1990s, the Cape Roberts Project represented the culmination of the previous two decades of sea-ice drilling experience, brought together through the planning efforts of its International Steering Committee, variously chaired by Maria Bianca Cita (Italy), Fred Davey (NZ), Franz Tessensohn (Germany), Mike Thomson (UK), Jaap van der Meer (Holland), Peter Webb (USA) and Ken Woolfe (Australia), and the Operations Management Group chaired by Gillian Wratt (NZ). Finally, I wish to acknowledge the work of the ACE community, led by Martin Siegert and Rob Dunbar, and within that the ANDRILL Steering Committee (the next generation) – Dave Harwood and Ross Powell (USA), Tim Naish and Gary Wilson (NZ), Fabio Florindo and Franco Talarico (Italy) and Frank Niessen and Gerhard Kuhn (Germany) – in ﬁlling in key gaps in Antarctic climate history in the McMurdo Sound area, and hopefully beyond in the future. Reviews by Peter Webb, Colin Summerhayes and Jane Francis on an earlier draft were very helpful. Preparation of this chapter was supported by NZ Foundation for Research, Science and Technology, Grant No. C05X0410 ANDRILL.

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Developments in Earth & Environmental Sciences, 8 F. Florindo and M. Siegert (Editors) r 2009 Elsevier B.V. All rights reserved DOI 10.1016/S1571-9197(08)00004-9

Chapter 4

Circulation and Water Masses of the Southern Ocean: A Review Lionel Carter1,, I. N. McCave2 and Michael J. M. Williams3 1

Antarctic Research Centre, Victoria University, P.O. Box 600, Wellington, New Zealand 2 Department of Earth Sciences, University of Cambridge, Downing Street, Cambridge CB2 3EQ, UK 3 National Institute of Water and Atmospheric Research, P.O. Box 14901, Wellington, New Zealand

ABSTRACT The Southern Ocean is a major component of Earth’s ocean and climate. Its circulation is complex, with a zonal Antarctic Circumpolar Current (ACC) interacting with a meridional thermohaline circulation. The ACC is a highly variable, deep-reaching eastward ﬂow driven mainly by the westerly winds. It is the longest (24,000 km), largest (transport 137–147.106 m3 s1) and only current to connect the major oceans. The Ekman component of the westerly winds also drives surface waters north. Near the ACC’s northern limit, these waters sink to form Subantarctic Mode and Antarctic Intermediate waters, which continue north at depths oB1,400 m. Interacting with the ACC is the density-forced thermohaline circulation. Super cooling and increased salinity of shelf waters off the Weddell, Wilkes Land and Ross coasts cause these waters to sink and ﬂow equatorwards. The densest component, Antarctic Bottom Water, is captured in deep basins around Antarctica. Less dense water is entrained by the ACC and mixed with deep water moving south from the Atlantic, Indian and Paciﬁc oceans. The resultant Lower Circumpolar Deep Water is tapped off by deep western boundary currents that enter the three oceans at depths WB2,000 m. These northward inﬂows, with a total volume transport of B55.106 m3 s1, disperse Antarctic and

Corresponding author. Tel.: þ64 4 463 6475; Fax: þ64 4 463 5186;

E-mail: [email protected] (L. Carter).

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northern-sourced waters throughout the world ocean. Other circulation elements are the deep-reaching, cyclonic Weddell, Ross and unnamed gyres located south of the ACC. Further south again are the westward Antarctic Slope and Coastal currents that pass along the Antarctic continental margin under easterly polar winds.

4.1. Introduction The Southern Ocean has a profound inﬂuence on the world’s ocean and climate. Cold, dense water sinks to abyssal depths around the margins of Antarctica and migrates northwards into the Atlantic, Indian and Paciﬁc oceans via deep western boundary currents (Fig. 4.1; Stommel, 1958; Warren, 1981). As succinctly noted by Warren (1971), ‘y water from the Antarctic is largely responsible for keeping the rest of the deep sea cold’. Through a process of slow upwelling, these deep cold waters rise to the upper ocean. There, they contribute to the warm surface circulation that extends west from the Paciﬁc and Indian Oceans into the Atlantic where the warm, saline water moves north. Approaching high northern latitudes, the water cools and sinks to form North Atlantic Deep Water (NADW), which migrates south, sandwiched between northbound Antarctic Intermediate Water (AAIW) above and Antarctic Bottom Water (AABW)/Lower Circumpolar Deep Water (LCDW) below (Fig. 4.2). En route, NADW mixes with other waters and eventually rises at the Antarctic continental margin. Thus, one cycle of the global thermohaline circulation (THC) – a major regulator of Earth’s ocean and climate – is completed and another cycle begins (e.g. Broecker, 1991; Schmitz, 1995; Rahmstorf, 2002). This powerful and far-reaching inﬂuence of Antarctica and the surrounding Southern Ocean largely reﬂects; (i) the strong buoyancy-driven and meteorologically forced circulations, and (ii) their direct access to the major ocean basins via the Antarctic Circumpolar Current (ACC) and its offshoots, the deep western boundary currents (Fig. 4.1; Moore et al., 1999; Orsi et al., 1999; Rintoul et al., 2001). In this brief synopsis we can only provide a ﬂavour of over 70 years of oceanographic research in the Southern Ocean. Thus, we refer the reader to the reference list for a more detailed insight into the workings of this region. We present the basic elements under two sections: (1) Section 4.2 examines the main water masses, focusing on their properties and the mechanisms that control their distribution, and (2) Section 4.3 reviews the structure and dynamics of the world’s largest ocean current, the ACC, together with that of the subpolar gyres and

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180° 150°W

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Figure 4.1: The main Oceanographic elements of the Southern Ocean including: (i) the ACC contained by the Subantarctic Front (SAF) and southern limit of UCDW or southern boundary (SB); (ii) the Ross, Weddell and unnamed subpolar gyres; and (iii) the main exit points of deep western boundary currents from the Southern Ocean (blue arrows). The general path for the ACC is from Orsi et al. (1995) with modiﬁcations based on Heath (1985) and Morris et al. (2001). Bathymetric elevations are Annotated as R., ridge; K. Pl., Kerguelen Plateau; and F. Pl., Falkland Plateau. The base chart is Modiﬁed from Orsi and Whitworth (2005). currents residing south of the ACC, and the deep THC. The chapter ends with a discussion in Section 4.4 of the present debate regarding the Southern Ocean’s response to a rapidly warming climate (e.g. Gille, 2002; Curry et al., 2003; Jacobs, 2004; IPCC, 2007).

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ACC 0

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Figure 4.2: Schematic section of the main water masses and their meridional transport as compiled from Whitworth (1988); Orsi et al. (1995); Speer et al. (2000) and Rintoul et al. (2001). Water masses are SAMW, Subantarctic Mode Water; AAIW, Antarctic Intermediate Water; UCDW, Upper Circumpolar Deep Water; LCDW, Lower Circumpolar Deep Water; NADW, North Atlantic Deep Water; AABW, ‘true’ Antarctic Bottom Water (gnW28.27 kg m3). Frontal systems are ASF, Antarctic Slope Front; SB, Southern Boundary of the ACC; SF, Southern Front; PF, Polar Front (formerly the Antarctic Convergence), SAF, Subantarctic Front; STF, Subtropical Front (formerly Subtropical Convergence). The ﬂow of ACC is directed towards the reader.

4.2. Water Mass Formation and Dispersal 4.2.1. Surface Ocean A series of ocean fronts – narrow, variable bands deﬁned by abrupt changes in water properties, in particular, temperature and salinity – divide the surface waters of the Southern Ocean into several zones (e.g. Gordon, 1975; Deacon, 1982; Whitworth, 1988). Early studies identiﬁed (from south to north) the Polar, Subantarctic and Subtropical fronts (Fig. 4.2). More recent hydrographic transects, especially those carried out during the World Ocean Circulation Experiment (WOCE), have revealed additional boundaries located south of the Polar Front, and termed the ‘southern’ and ‘southern

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boundary’ fronts (Figs. 4.2 and 4.3; Orsi et al., 1995; Orsi and Whitworth, 2005). Furthermore, these detailed and sometimes repeated transects, along with satellite-borne observations of ocean height, temperature and drifter tracks, reveal the complex and dynamic character of the frontal systems 180° 150°W

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Figure 4.3: Location of the principal ocean frontal systems in the Southern Ocean (based on Orsi et al., 1995, but modiﬁed for the New Zealand region according to Carter et al., 1998 and Morris et al., 2001). Repeated hydrographic transects, satellite observations and drifting ﬂoats reveal the frontal systems as dynamic features with marked temporal and spatial variability but generally within the constraints imposed by the ocean ﬂoor topography (Moore et al., 1999). P14 and S2 are locations of WOCE hydrographic transects portrayed in Figs. 4.4 and 4.5. Madg., Madagascar Basin; Mozb., Mozambique Basin. Names of fronts are given in Fig. 4.2. The base chart is modiﬁed from Orsi and Whitworth (2005).

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(Hofmann, 1985; Davis, 1998; Moore et al., 1999; Rintoul et al., 2001; Kostianoy et al., 2004; Sokolov and Rintoul, 2007). While cognizant of these complexities, the main fronts can still be used to deﬁne the distribution of three major surface waters characterised mainly by their potential temperature (y), salinity (S) and oxygen content (see hydrographic charts in Orsi and Whitworth, 2005). (1) Near-freezing and relatively fresh Antarctic Surface Water (AASW) forms a layer about 100 m thick that extends from the Antarctic continental shelf to the Polar Front, commonly deﬁned as the northernmost extent of the subsurface temperature minimum (Belkin and Gordon, 1996; Figs. 4.4 and 4.5). AASW temperatures are typically o01C, but may rise to 2.51C near the Front or where warm, deep water approaches the surface (Gordon, 1975; Deacon, 1982). Salinity (S) varies regionally with highest values of SW34.3 psu found in the Ross and Weddell seas, whereas AASW elsewhere around Antarctica commonly has So34.0 psu. (2) Between the Polar and Subantarctic fronts resides surface water that is transitional between AASW and Subantarctic Surface Water (SASW). The structure is complex in response to mixing and interleaving of AASW as it sinks near the Polar Front (e.g. Gordon, 1975; Rintoul et al., 2001). Thus, properties are variable, but generally S is B34.0–34.4 psu and y is 3–81C (Fig. 4.5). (3) SASW occurs north of the Subantarctic Front and encompasses water that usually warms northwards from B61C to 121C. Salinity is typically W34.3 psu except in the SE Paciﬁc and Drake Passage where values decline to o34.16 psu. Like its more southern counterpart, SASW may be affected by vertical mixing as surface waters subduct and mix (e.g. Morris et al., 2001). The northern limit of subantarctic waters is the Subtropical Front where temperatures increase sharply by 4–51C and salinity by 0.5 psu (Fig. 4.3; Deacon, 1982). Subtropical surface water prevails north of the Subtropical Front. 4.2.2. Subantarctic Mode Water and Antarctic Intermediate Water Isopycnals – surfaces of constant density – of near-surface to deep waters in the Southern Ocean rise up in a step-like proﬁle towards Antarctica. Close to ocean fronts, isopycnals may outcrop indicating either the rise of deep waters to the ocean surface or the descent of surface waters (Figs. 4.2, 4.4 and 4.5). In the case of the latter, winter cooling and mixing of the surface waters on the northern side of the Subantarctic Front forms Subantarctic Mode Water (SAMW) (McCartney, 1977; Morris et al., 2001; Rintoul et al., 2001). This well mixed, ventilated water descends northwards to B500 m depth along much of the front (Fig. 4.2). AAIW also descends from the surface, passing

Figure 4.4: Sections of potential temperature (A), salinity (B) and neutral density (C) from WOCE Line P14 across a major constriction in the ACC between the Ross Sea and New Zealand (see Fig. 4.3 for location). Isolines at the Antarctic margin indicate descent of dense shelf waters, which may be mixed with NADW-inﬂuenced, LCDW as suggested by the salinity ﬁeld. The resultant AABW (cf. Fig. 4.5) is contained within the SW Paciﬁc Basin. At the surface, north of the polar front, low salinity AAIW descends northwards (Fig. 4.4B). Hydrographic proﬁles were derived from the WOCE Southern Ocean Atlas at http://woceatlas.tamu.edu/. Names of fronts are given in Fig. 4.2.

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Figure 4.4: (Continued). under SAMW to reach a maximum depth of B1,400 m (Figs. 4.2 and 4.4). AAIW is identiﬁed by a salinity minimum (34.3–34.5 psu) and temperatures of B3–71C. However, the processes driving AAIW formation are unclear. Formation may be related to wind-forced or density-driven sinking of cold AASW and indeed there appears to be continuity between the AASW and AAIW salinity ﬁelds (Figs. 4.4 and 4.5). However, McCartney (1977) suggested AAIW may evolve at least in part from dense SAMW. Whatever the origin, compared to the widespread formation of SAMW, new AAIW presently appears to form mainly in the SW Atlantic and SE Paciﬁc. From these sites AAIW circulates the oceans in anticyclonic subtropical gyres that extend north towards and locally beyond the equator before returning south as ‘old’ AAIW, which is transported within western boundary currents (Rintoul et al., 2001; Ridgway and Dunn, 2007). 4.2.3. Circumpolar Deep Water The most voluminous water mass in the Southern Ocean is Circumpolar Deep Water (CDW). It extends from B1,400 m to W3,500 m depth, but it

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Figure 4.5: Hydrographic sections of the salinity (A) and neutral density (B) ﬁelds from WOCE Line S2 in the Atlantic sector of the Southern Ocean. Of note is (i) the step-wise ascent of saline, NADW-rich LCDW towards Antarctica where it is capped by fresh, cold AASW in the south and subducting AAIW in the north and (ii) the containment of a large volume of classic Antarctic Bottom Water (AABW; gnW28.27 kg m3) within the Weddell Basin (cf. Fig. 4.4). Hydrographic proﬁles are derived from the WOCE Southern Ocean Atlas at http://woceatlas.tamu.edu/. Names of fronts are given in Fig. 4.2.

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rises to meet AASW or even outcrop along the Antarctic continental margin (Figs. 4.2, 4.4 and 4.5). CDW has two basic types: (1) Upper Circumpolar Deep Water (UCDW) is identiﬁed by the oxygen minimum and elevated nutrient concentrations, and has an open ocean depth range of B1,400– 2,500 m, and (2) Lower Circumpolar Deep Water (LCDW) whose signature is the salinity maximum (34.70–34.75 psu) (Gordon, 1975; Orsi et al., 1995). This maximum reﬂects the input of NADW that has migrated into the Southern Ocean (Orsi et al., 1995; Rintoul et al., 2001). Upon meeting the ACC, NADW is entrained and transported east around the Antarctic continent, all the while mixing with waters from the Indian and Paciﬁc oceans plus dense waters from Antarctica to form LCDW. Despite the vigorous mixing, the high-salinity signature of NADW is retained (Reid and Lynn, 1971; Gordon, 1975). At several locations around Antarctica, LCDW rises at the continental slope where mixing with super-cold shelf water renews not only the deep circulation of LCDW but also generates Antarctic Bottom Water (AABW), the deepest water mass in the Southern Ocean (Figs. 4.2, 4.4 and 4.5; Foster and Carmack, 1976; Jacobs et al., 1985; Orsi et al., 1999). LCDW is carried equatorwards in all three major oceans by deep western boundary currents (see Section 4.3.3 and Schmitz, 1995; Hogg, 2001). According to Rintoul et al. (2001) mixing with fresher waters, together with the biological depletion of oxygen, slowly modify LCDW into a less dense, lower oxygenated variant that returns south as UCDW. In some WOCE sections, such as P15 in the Indian Ocean, oxygen depletion of UCDW may be inﬂuenced by the direct injection of nutrient-enriched deep waters from the North Indian and North Paciﬁc oceans. 4.2.4. Antarctic Bottom Water In their analysis of abyssal water masses, Mantyla and Reid (1983) drew attention to the often inappropriate use of the term Antarctic Bottom Water, which has tended to be used generically for any southern-sourced bottom water. They demonstrated that true AABW did not extend far from Antarctica before mixing with other waters (see also Orsi et al., 1999). This is particularly true for the deep western boundary currents in which AABW is mixed with CDW derived from the ACC. Thus, at 301N in the NW Atlantic, Amos et al. (1971) recorded o20% AABW near the seabed. To better characterise AABW and thereby improve assessments of its dispersal and contribution to bottom waters worldwide, Orsi et al. (1999) deﬁned AABW by its neutral density (gn) whereby gnW28.27 kg m3. Such dense waters are conﬁned mainly to the deep (down to B6,000 m),

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circum-Antarctic Basins that include the Argentine and Brazil basins (SW Atlantic), Mozambique, Crozet and Australian-Antarctic basins (Indian) and the SE Paciﬁc Basin (Figs. 4.1, 4.4–4.6). Less dense, southern bottom water with 28.18ogno28.27 kg m3, is not conﬁned to the circum-Antarctic basins, but instead spreads out from the deep levels of the northern edge of the ACC into all major oceans (Orsi et al., 1999; also see Section 4.3.3). AABW density is determined by a combination of different sources that produce regionally distinct waters (Mantyla and Reid, 1983; Orsi et al., 1999; Jacobs, 2004). The freshest and coldest (Sr34.64 psu; yr11C) bottom water occurs in the Weddell Sea, whereas the SE Paciﬁc Basin has the most saline and least cold (SZ34.72 psu; 0.6oyo0.31C) bottom water. The Australian-Antarctic Basin contains water with properties intermediate between the two end members. Traditionally, the Weddell Sea was regarded as the prime source of AABW, but recently two other sources have come to the fore. The Weddell Sea’s contribution is now regarded as B50%, with the Wilkes Land margin including the Ade´lie coast (Rintoul, 1998) contributing B30% mainly to the Indian Ocean, and the Ross Sea producing B20% AABW that is destined primarily for the SE Paciﬁc Basin (Jacobs, 2004). Equally varied are the modes of bottom water formation (Jacobs, 2004). Foster and Carmack (1976) invoked the formation of highly saline shelf water (HSSW) by brine rejection from sea ice. However, salt-driven increases in density may also be inﬂuenced by intrusions of NADW-bearing LCDW onto the upper slope and shelf (Toggweiler and Samuels, 1995). Super cold, Ice Shelf Water (ISW), formed by freezing and melting below ice shelves, can mix with HSSW and reach the outer shelf before ﬂowing down slope (Baines and Condie, 1998). Alternatively, the simple mixing of cold AASW and highsalinity LCDW, with or without ISW, may produce negatively buoyant waters at the shelf edge (e.g. Jacobs, 2004). Finally, dense waters may sink via convection chimneys and polynyas such as the well-documented but short-lived Weddell Sea polynya (Gordon, 1982). Rates of AABW formation, as estimated from hydrographic data, usually fall within a range of 5–15 Sv (1 Sv ¼ 106 m3 s1). Anthropogenic tracers, in particular chloroﬂuorocarbons, record a ﬂux of 8.1 Sv for AABW descending at the 2,500 m isobath off Antarctica (Orsi et al., 2001). This compares to 7.6 Sv of lower NADW ﬂowing out of the Nordic and Labrador seas at the N Atlantic source. Because AABW is colder than its northern counterpart, Antarctic overturning probably plays the dominant role in cooling the deep ocean. The time at which deep water circulates through the ocean is often quoted to be a millennium or more, for example 1,000 years between the N Atlantic and Southern Oceans and a further 1,000 years from the Southern

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Figure 4.6: Extent of dense Antarctic Bottom Water as identiﬁed by the neutral density ﬁeld at 3500 m (Orsi et al., 1999; Orsi and Whitworth, 2005). With the exception of the SE Atlantic, where AABW extends well north via the deep Argentine and Brazil basins, the remaining AABW is captured in circum-Antarctic basins where further northwards dispersal is inhibited by oceanic ridge systems shown in white. Basins annotated as Arg. B., Argentine Basin, which extends north into the Brazil basin (not shown on chart); W.B., Weddell Basin; E.B., Enderby Basin; C.B., Crozet Basin; Aust. A.B., Australian-Antarctic Basin, and SE P.B., SE Paciﬁc Basin. Chart is generated from the WOCE Southern Ocean Atlas at http://woceatlas.tamu.edu/.

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Ocean to the North Paciﬁc. However, in a re-analysis of radiocarbon dated ocean waters, Matsumoto (2007) indicates much shorter circulation ages thereby supporting but reﬁning earlier radiocarbon-based studies (Stuiver et al., 1983). Thus, the circulation age for the Southern Ocean below 1,500 m is B300 14C years with a similar age for the Atlantic. For the Paciﬁc, the basin circulation age is B900 14C years.

4.3. Ocean Circulation 4.3.1. Antarctic Circumpolar Current The Southern Ocean circulation is dominated by the ACC, a current system that is rightly described by superlatives. It is the only current to connect the major ocean basins and hence plays a prominent role in the global distribution of heat, salt and gases (Fig. 4.1). It is the longest current with an estimated pathway of B24,000 km (Whitworth, 1988). Finally, the ACC is the largest major current in terms of volume transport with a mean of 136.777.8 Sv as measured across Drake Passage (Cunningham et al., 2003). The ACC was originally termed the West Wind Drift (Deacon, 1937) in recognition of its forcing by middle-to-high latitude westerly winds (Orsi et al., 1995; Whitworth et al., 1998). However, use of the term wind drift masks the role played by the buoyancy-driven component of the circulation (see Rintoul et al., 2001). Complexities aside, the net result is an eastward current system that extends from the ocean surface to bottom, its path guided by submarine topography (Fig. 4.1; Gordon et al., 1978; Orsi et al., 1995). For much of its passage, the ACC ﬂows along the ﬂanks of midoceanic ridges except within major gaps in the Paciﬁc and Indian ridge systems where the current shifts poleward (Fig. 4.1). Large submarine plateaux also exert an inﬂuence. The ACC widens to the north and south as it passes around the Kerguelen Plateau, whereas the Campbell and Falkland plateaux form constrictions (Fig. 4.1; Whitworth and Peterson, 1985; Morris et al., 2001). Interestingly, this interaction with the ocean ﬂoor was inferred as early as the 1950s. Estimates for a purely wind-driven ACC yielded transports that were excessive compared to observations. Thus, it was concluded that the wind stress was partly balanced by bottom stress (see Whitworth, 1988; Rintoul et al., 2001). The passage of westerly winds over the ACC also induces an Ekman drift to the north – a process that probably plays a role in the subduction and transport of mode and intermediate waters. This equatorward ﬂow is compensated at depth by the southward

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transport and eventual upwelling of CDW thus contributing to the THC (Wyrtki, 1961; Toggweiler and Samuels, 1995). Rather than a uniform ﬂow, the ACC is a system of deep-reaching zonal jets that separates zones of relatively quiet water. The jets are marked by the circumpolar fronts (see Section 4.2 Surface Ocean) with the northern and southern boundaries of the ACC deﬁned, respectively, by the Subantarctic and Southern Boundary fronts (Figs. 4.1–4.3). Both eddy-resolving models and satellite observations highlight the complex ﬂow of the frontal jets (Morrow et al., 1992; Gille, 1994; Carter and Wilkin, 1999). Meanders, eddies and intricate branches are well shown especially where the ﬂow is constricted as off Campbell Plateau and within Drake Passage (Nowlin and Klinck, 1986; Morris et al., 2001; Cunningham et al., 2003). Most ACC transport takes place within the fronts, but their complex ﬂow patterns and different criteria for estimating transport have led to a wide range of values for the entire ACC. Nevertheless, closely spaced and longterm monitoring sites have improved estimates of transport. In Drake Passage the mean transport is 136.777.8 Sv (Cunningham et al., 2003) compared to 147710 Sv between Australia and Antarctica (Rintoul and Sokolov, 2001). Much of the transport in the Australasian reach of the ACC occurs within the Subantarctic Front, which has a mean of 10577 Sv off Australia (Rintoul and Sokolov, 2001) and B90 Sv off southern New Zealand (Morris et al., 2001). However, ACC transport can be highly variable. Time series from Drake Passage record variations at several time scales (Whitworth and Peterson, 1985). Short-term ﬂuctuations, related to 14-day solar and lunar tides, are superimposed on longer-period ﬂuctuations of B1 year and longer that can lead to changes in transport of B30–40 Sv within weeks. The interaction of the ACC with the topography and southerly extensions of western boundary currents generates eddies that play important roles in the transfer of heat and momentum (Bryden and Heath, 1985; Morrow et al., 1992; Rintoul et al., 2001). Off SE New Zealand, for example interception of the ACC by the South Tasman Rise and Macquarie Ridge spawns bottomreaching eddies that migrate NE along the steep margin of Campbell Plateau (Boyer and Guala, 1972; Gordon, 1972). Both cyclonic and anticyclonic features have been observed from current meter and satellite data, which suggest a frequency of occurrence of B9 eddies annually (Stanton and Morris, 2004). Modelled eddy kinetic energy, veriﬁed by current meter data and ocean ﬂoor sedimentary evidence, attest to the power of these perturbations, which are likely to be the cause of abyssal benthic storms (Hollister and McCave, 1984; Carter and Wilkin, 1999). Seabed topography also encourages intense mixing within the ACC. The Scotia Sea and

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potentially other areas of marked seabed relief below the ACC, are zones of the highest turbulent mixing in the ocean and result in rapid upwelling that may locally short-circuit the classic meridional overturning as portrayed in Fig. 4.2 (Garabato et al., 2004, 2007).

4.3.2. Subpolar Circulation 4.3.2.1. Major gyres A key element of the circulation, south of the ACC, is three large, deepreaching cyclonic gyres that extend from the ACC to the Antarctic continental margin (Fig. 4.1). The better known are the Weddell and Ross gyres that occupy ridge-bounded sectors of the Weddell and Ross seas (e.g. Orsi et al., 1993; Jacobs et al., 2002). A third and as yet unnamed gyre was suggested to occur south of Kerguelen Plateau (Bindoff et al., 2000), and has now been conﬁrmed by McCartney and Donohue (2007). As documented for the Weddell Gyre (Orsi et al., 1993), it is likely that all three gyres transport salt and heat from the ACC to the Antarctic continental margin where deep and bottom waters are produced (e.g. Jacobs, 2004). Furthermore, McCartney and Donohue (2007) suggest a strong connectivity between the three gyres with a westward ﬂow along their southern limbs and an eastward ﬂow joining their northern limbs. This latter ﬂow is just south of another eastward ﬂow, this time associated with an anticyclonic supergyre covering most of the S Paciﬁc and S Indian oceans (Ridgway and Dunn, 2007). This eastward limb of the supergyre appears to reside between the Subtropical and Subantarctic fronts. As a result, it may entrain SAMW, formed in the vicinity of the Subantarctic Front and help distribute it through the ocean basins as suggested by Rintoul et al. (2001). Of these cyclonic circulations, the Weddell Gyre is the largest, extending from B501W to between 201–301E (Fig. 4.1; Orsi et al., 1993). At the surface it has the form of a NE–SW aligned, elongated gyre whereas at depth it comprises two cyclonic cells located to the east and west of 151W. Basically, the Weddell Gyre occupies the W4,000 m deep re-entrant formed by Antarctica and the ridge systems that extend eastward from near the tip of the Antarctic Peninsula (Fig. 4.1). While such a location implies containment, the northern limb of the gyre overlaps the ridges to interact with the ACC. CDW entrained from the ACC is moved within the gyre and can eventually mix with cold shelf waters to form Weddell Sea Bottom Water, the local variety of AABW and the densest water in the Southern Ocean (Foster and Carmack, 1976).

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The Ross Gyre extends from 1601E to 1401W, and is largely conﬁned to the W4,000 m deep western reach of the SE Paciﬁc Basin (Fig. 4.1). Like its Weddell Sea counterpart, the Ross Gyre is a deep-reaching feature that entrains CDW to make it available for mixing with shelf and slope waters. Situated to the south and southwest of Australia, the unnamed gyre appears to favour the southern part of the Australian-Antarctic Basin, most of which is W4,000 m deep (Fig. 4.1). Indeed, the volume transport ﬁeld is compressed against the Antarctic continental margin and the associated westward slope current (Bindoff et al., 2000; McCartney and Donohue, 2007). Transport along the eastward-ﬂowing northern limb of the gyre is estimated at 35 Sv. However, the amount of transport along the westmoving, southern limb is unclear due to merging with the slope current, the combined ﬂows reaching 76 Sv (McCartney and Donohue, 2007).

4.3.2.2. Antarctic slope and coastal currents As summarised by Heywood et al. (2004) much of the Antarctic margin is bathed by two westward currents. One is associated with the Antarctic Slope Front that constitutes the boundary between fresh, cold Antarctic shelf waters and less cold, saline CDW (Jacobs, 1991; Whitworth et al., 1998). Deacon (1937) regarded the frontal ﬂow as a consequence of the prevailing easterly winds and coined the name, East Wind Drift. On the basis of classical theory, he reasoned that polar easterly winds produced an onshore Ekman transport with the resultant generation of a westward geostrophic current below the wind-mixed layer (Whitworth et al., 1998; Bindoff et al., 2000). The second signiﬁcant feature is the Antarctic Coastal Current, which forms a narrow, rapid ﬂow across broad sections of the continental shelf, for example, in the SW Weddell Sea. However, where the shelf is narrow the coastal current is difﬁcult to distinguish from ﬂows associated with the Antarctic Slope Front and the southern limbs of the subpolar gyres where they approach the continental margin. 4.3.3. Thermohaline Circulation Wunsch (2002) drew attention to the imprecise meaning of the term, THC. Following Jacobs (2004), we prefer the broad deﬁnition of Schmitz (1995) whereby the THC is the ‘y buoyancy-driven ﬂow ﬁeld associated with water cooled (or heated) by contact with cold (or warm) air, or modiﬁed by sources and sinks of cold water. May also include ﬂows whose characteristics are

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signiﬁcantly altered by upwelling and/or mixing. Water sinking at high latitudes tends to return equatorwards in relatively strong, narrow currents called DWBC’. The high latitude sinks noted are primarily the North Atlantic and Antarctica (Stommel, 1958; Warren, 1981). As described earlier, dense water formed mainly over shelf areas of the Weddell and Ross seas, and along the Wilkes Land coast, sinks and descends down the continental slope. Descent may take place in several ways according to Baines and Condie (1998). For prominent deep-water formation areas, dense water may descend as broad sheets or plumes. For weaker sources, the resultant outﬂows may geostrophically descend the slope to a level where the layer thins and viscous drainage prevails. Descending ﬂows may be disrupted by eddies into discrete parcels that overall move down slope. Further complexities are introduced by the slope topography, for example, submarine canyons and channels can contain and steer ﬂows whose density may increase through entrainment of sediment. Whatever the mode of descent, the initial dispersal of AABW is westward as outlined by the distribution of chloroﬂuorocarbons (Orsi et al., 1999). Those tracers show that the densest AABW usually follows a cyclonic path presumably under the inﬂuence of the subpolar gyres and basin topography (Fig. 4.1). However, the Weddell– Enderby Basin experiences limited outﬂow at its northern rim where least dense AABW (gnB28.28 kg m3) passes north into adjoining basins of the Atlantic and Indian oceans (Fig. 4.6; Orsi et al., 1999). In contrast, the main inﬂows to the adjoining oceans are via deep western boundary currents, which carry mainly LCDW from the northern boundaries of the ACC (Mantyla and Reid, 1983; Schmitz, 1995).

4.3.3.1. Atlantic DWBC inﬂow Weddell Sea Bottom Water, a local variety of AABW, together with Weddell Sea Deep Water and CDW from the Drake Passage, are carried by a northbound DWBC into the Atlantic Ocean below southward-moving NADW (Warren, 1981; Schmitz, 1995). The DWBC is typically found in water depths exceeding 3,500–4,000 m along the western boundary presented by the continental margin off South America. However, the current pathway is interrupted by a succession of deep basins including (from south to north) Georgia, Argentine and Brazil basins. As noted earlier, densest AABW is captured within the deep basins leaving less dense waters to move north via gaps and channels through the inter-basin ridges. Even so, not all the deep water escapes and some recirculates within the basins themselves (Hogg and Johns, 1995; Hogg, 2001). For example, around 6.9 Sv ﬂows from the

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Argentine Basin into the Brazil Basin through Vema and Hunter channels, but only 3.2 Sv leaves the Brazil Basin, leaving 3.7 Sv to recirculate.

4.3.3.2. Indian DWBC inﬂow The inﬂow of AABW into the western Indian Ocean is via Crozet and Mozambique basins (the latter being a dead end), which largely constrain AABW to south of 301S and 341S, respectively (Orsi et al., 1999). LCDW from the Crozet Basin leaks northwards through fractures zones in the SW Indian Ridge (Warren, 1974; Johnson et al., 1991; Mantyla and Reid, 1995). Deep transport is via a northbound DWBC, which like its Atlantic counterpart, has a complex pathway dictated by multiple basins and ridges (Warren, 1981; McCave et al., 2005). Basins encourage recirculation of deep waters and together with mixing, account for a northward dissipation of the DWBC. For instance, LCDW below B3,800 m has a northwards transport of 3.8 Sv through Madagascar Basin, but only 1.7 Sv exits north into the Somali Basin (Johnson et al., 1998). Crozet Basin, in the western Indian Ocean, is not the only gateway for DWBCs into the Indian Ocean. The eastern Indian Ocean is also connected to the Southern Ocean thereby allowing the northward intrusion of two other boundary currents; one along the eastern side of SE Indian Ridge and the other along the eastern ﬂank of Ninetyeast Ridge (Warren, 1981; Toole and Warren 1993; Reid, 2005). Again, the deep water carried north is LCDW, with the denser AABW retained in Australian-Antarctic Basin (Orsi et al., 1999). When extended to the 2,000 dbar reference level, the combined northward transport of the three DWBCs into the Indian Ocean is B27 Sv (Toole and Warren, 1993).

4.3.3.3. Paciﬁc DWBC inﬂow The general pathway of the DWBC into the Paciﬁc was ﬁrst outlined in the classic model of the global abyssal circulation by Stommel (1958) and Stommel and Arons (1960), and was later conﬁrmed by the hydrographic sections of Warren (1971, 1973). Because the Tasman Basin is essentially closed at its northern end, the main inﬂow is off southernmost New Zealand where the DWBC enters in concert with the ACC (Fig. 4.7; Carter and McCave, 1997; Carter and Wilkin, 1999). Initially, the combined inﬂow intercepts Macquarie Ridge to form meanders and eddies although some current ﬁlaments pass through narrow gaps in the ridge (Boyer and Guala,

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Figure 4.7: A schematic portrayal of the THC, which is a series of loosely linked, recirculation systems that transport, heat, salt, nutrients and ventilating gases through the world ocean. Only the main elements of the THC are shown, for example the Indian Ocean has three, deep northward inﬂows that include from W to E, the margin off eastern Africa (shown), and the eastern sides of the SE Indian Ridge and Ninetyeast Ridge (not shown) (image modiﬁed from Manighetti, 2001).

1972; Gordon, 1972). This perturbed combined ﬂow continues northeast along the 3,000–3,500 m high ﬂanks of Campbell Plateau to around 491S where the ACC veers east leaving the DWBC to continue northwards into the Paciﬁc Ocean (Fig. 4.7). It eventually departs the Southern Ocean off Chatham Rise between 441S and 421S (McCave and Carter, 1997). There, Warren (1973) observed a volume transport of B20 Sv, the largest for a single DWBC (Schmitz, 1995).

4.4. Oceanographic Variability and Change The present phase of climate change has drawn considerable attention to the behaviour of Antarctica and the Southern Ocean (e.g. Gille, 2002; Jacobs et al., 2002; Walther et al., 2002; Cook et al., 2005). However, conﬁdent identiﬁcation of any effect, especially in the oceans, has been

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hindered by; (i) sampling bias, (ii) the short history of observations, which typically only encompass the last 50–60 years, (iii) a multiplicity of forcing mechanisms, some with and some without clear connections to a warming climate and (iv) marked variability at a range of temporal and spatial scales (e.g. Jacobs and Giulivi, 1998; Orsi et al., 2001; Vaughan et al., 2001). Of special note is the Southern Annular Mode (SAM), which appears to dominate inter-annual to centennial variability in the Southern Ocean (Hall and Visbeck, 2002; Lovenduski and Gruber, 2005; Lovenduski et al., 2007). SAM is essentially a zone of climate variability that encircles the South Pole and strongly inﬂuences zonal winds, sea ice formation and oceanic circulation. When in a positive phase, SAM is typically associated with enhanced westerly winds over the ACC and weakened winds further north. This favours a strengthening of the ACC, a reinforcement of the northward Ekman drift and subduction of surface waters, and a northward expansion of sea ice. To compensate, there is enhanced poleward transport and rising of deep water at the Antarctic continental margin. Under a negative SAM, windiness and storminess appear to migrate to mid-latitudes thus weakening both zonal and meridional ocean transport. Despite uncertainties associated with the aforementioned limitations, it is nonetheless important to examine and evaluate recent changes in Antarctica and the Southern Ocean in light of their actual or potential inﬂuence on the global ocean and climate. Climatic trends for the past 50 years have been identiﬁed by Turner et al. (2005) using records from 19 Antarctic meteorological stations. The results emphasise the marked geographic variability of the continental climate (e.g. Vaughan et al., 2001) as well as its temporal variability at interdecadal scales. The Antarctic Peninsula has warmed at a statistically signiﬁcant rate of þ0.561C/decade from 1951 to 2000. The next largest warming trend outside of the Antarctic Peninsula is in the western Ross Sea (Scott Base) with a rise of þ0.291C/decade, although the rise is not statistically signiﬁcant. Elsewhere, signiﬁcant trends are unclear. Annual temperature trends for coastal and interior sites on the East Antarctic Ice Sheet suggest a slight cooling. However, all but one site exhibited warmer winters. Notwithstanding its temporal and spatial variability, the upper Southern Ocean has warmed between 1955 and 2003 in concert with the world ocean (Levitus et al., 2005). Off the Antarctic Peninsula, the pronounced atmospheric warming has been accompanied by an equally marked warming of the surface ocean with summer temperatures increasing by 1.21C over the latter half of the twentieth century (Meredith and King, 2005). At water depths of 700–1,100 m, Gille (2002) reported an average 0.171C rise since 1950, which accounts for about two-thirds of the total increase in heat

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content in the ocean from 0 to 3,000 m depth (IPCC, 2007). Much of the warming is concentrated within the Subantarctic Front of the ACC. This raises the possibility that the warming at depth may result from the sinking of atmospherically warmed SAMW. Gille (2002) further suggests that warming may also be related to a general southward displacement of the ACC. This would be consistent with a suite of model simulations that suggest warming of the Southern Ocean will be accompanied by a southward shift of zonal westerly winds together with a narrowing and intensiﬁcation of the ACC (IPCC, 2007 and references therein). Coherent historical trends are also evident for salinity. Subpolar regions have generally become fresher between 1955 and 1998 in contrast to subtropical and tropical regions, which display increased salinity with the exception of the central Paciﬁc Ocean (Curry et al., 2003; Boyer et al., 2005). Such freshening of the upper Southern Ocean may be responsible for a reduction in the salinity of AAIW (Wong et al., 1999; Curry et al., 2003). Reduced salinity implies increased freshwater input that may result from one or more of the following causes: (i) greater net precipitation; (ii) changes in the extent of sea ice – a major contributor to winter salinity through brine rejection; (iii) increased melting of ice shelves, ice sheets and glaciers (e.g. Cook et al., 2005), and (iv) changes in the oceanography, especially any reduced upwelling of saline LCDW (Wong et al., 1999; Jacobs et al., 2002; Curry et al., 2003). Nearer Antarctica, hydrographic records spanning over 40 years show local variability in salinity trends. The upper 50 m of the ocean, west of the Antarctic Peninsula, has become more saline although underlying waters have freshened slightly in line with the regional trend. While the more saline surface conditions appear to be out-of-step with the strong glacier retreat on the Peninsula (e.g. Cook et al., 2005), Meredith and King (2005) suggest more saline conditions are consistent with the reduced sea ice cover and the timing of the hydrographic measurements. With less sea ice production there is less freshening of the ocean in the summer when most of the hydrographic measurements are made. On the opposite side of the continent, at Law Dome, ice core records spanning a century and longer, identify a 20% loss of sea ice since 1950 although this trend is strongly overprinted with cyclical variations with an B11 year frequency (Curran et al., 2003). Reduced sea ice along with increased precipitation and melt water from the West Antarctic Ice Sheet have been cited as contributing to the observed freshening of surface waters associated with the Ross Sea Gyre (Jacobs et al., 2002). But like Law Dome, the trend is obscured by cycles, this time by 5–6- and 9-year oscillations in HSSW formation (Assmann and Timmerman, 2005).

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Salinity and temperature changes have the potential to affect bottom water production and the THC. Consequently, these changes have received considerable attention (IPCC, 2007 and references therein). At glacialinterglacial time scales, palaeoceanographic evidence reveals marked variations in the position and degree of convective overturning of the N Atlantic sector of the THC (Rahmstorf, 2002 and references therein). During interglacial periods, overturning is most active and reaches its most northerly extent. In contrast, glacial periods are likely to witness a southward shift and possible slow-down in overturning. Any slow-down may be compensated by a greater production of bottom water from Antarctica as suggested by grain size (Hall et al., 2001), magnetic properties (Venuti et al., 2007) and diatom proxies (Stickley et al., 2001). However, such conclusions are sometimes at odds with geochemical tracers such as d13C (e.g. Moy et al., 2006) that point to little change in the passage of NADW through the Indian and Paciﬁc sectors at least over recent glacial-interglacial cycles. At millennial time scales, abrupt changes such as those associated with Heinrich events, may stop N Atlantic overturning altogether as the density of the surface ocean is reduced by rapid inﬂuxes of melt water (Rahmstorf, 2002). Again, cessation of N Atlantic production may be compensated by enhanced Antarctic production. However, responses to the latest phase of climate/ocean warming are unclear. In the N Atlantic, which is the best observed deepwater source, long-term trends are equivocal due to decadal variability, a paucity of long-term observations and other factors (IPCC, 2007). A similar situation applies to Antarctica where estimations of bottom-water production are inconsistent in response to: (i) natural cycles; (ii) differences in the deﬁnitions and techniques to estimate production rates, and (iii) a bias towards summer observations (Jacobs, 2004). On the basis of chloroﬂuorocarbons and 14C data, which allow water masses to be traced at decadal to century scales respectively, Orsi et al. (2001) revealed no decline in bottom water production over the twentieth century as indicated earlier by Broecker et al. (1999). Nevertheless, the changes recorded in recent historical times cannot be ignored. The freshening of the Ross Gyre over the last 50 years (Jacobs et al., 2002) and an accompanying downstream freshening of AABW in the adjacent Australia-Antarctic Basin (Aoki et al., 2005) are consistent with increased freshwater input. On a larger scale, the historical salinity data of Curry et al. (2003) also reveals a freshening of deep and bottom water at Antarctic and N Atlantic sources. Simulations by 19 model runs under IPCC greenhouse gas scenario, A1B (rapid economic growth, world population peaks mid-century, new and efﬁcient energy technologies with reliance on a range of sources) point to an average 25% reduction in N Atlantic overturning by the year 2100 (IPCC, 2007).

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None of the runs point to a shut-down; rather they favour reductions in overturning of up to 50%. While the Southern Ocean sector has received less attention from modellers, simulations based on a warmer or fresher ocean may enhance or stabilise N Atlantic overturning (Saenko et al., 2003; Weaver et al., 2003). To further emphasise the complexity of north–south relationships, the projected strengthening of the Southern Hemisphere westerly winds will increase the northward Ekman transport of upper ocean. To compensate, the poleward ﬂow of deep water below 2,000–2,500 m depth, could be expected to strengthen and possibly stimulate the southward ﬂow of NADW (e.g. Toggweiler and Samuels, 1995; Toggweiler et al., 2006). Again, such a trend is overprinted with marked inter-annual variability. Because of the importance, size and complexity of the Southern Ocean, the incompleteness of observations, and its high variability at a range of temporal and spatial scales, it is critical to improve our knowledge of this ocean/climate system. To re-emphasise the introduction to this chapter, the Southern Ocean has a profound inﬂuence of the distribution of salt, heat and ventilating gases throughout global seas. At the same time it is also undergoing some of the most rapid environmental changes on Earth highlighted by the warming, glacial retreat and ice shelf collapse around the Antarctic Peninsula. Thus, to address the inevitable questions relating to impacts of a rapidly changing climate on the Southern Ocean we require a strong modelling effort, supported by multi-seasonal oceanographic and remotely sensed observations and high-resolution palaeoceanographic records of past warm extremes. While this may seem to be a well-worn message, it is still appropriate at a time of certain change and uncertain consequences.

ACKNOWLEDGEMENTS We are indebted to the World Ocean Circulation Experiment (WOCE) for permission to use their data for Figures 4.1, 4.3 to 4.6, which are attributed to Orsi, A. H., T. Whitworth III, Hydrographic Atlas of the World Ocean Circulation Experiment (WOCE). Volume 1: Southern Ocean (eds. M. Sparrow, P. Chapman and J. Gould), International WOCE Project Ofﬁce, Southampton, U.K., ISBN 0-904175-49-9, 2005. The chapter beneﬁted from the critiques by the external reviewers, Will Howard and Alejandro Orsi and their input is appreciated. Funding for L. Carter was provided by the Foundation for Research Science and Technology, contracts CO50410 and VICX0704.

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Whitworth III, T., Orsi, A. H., Kim, S.-J., & Nowlin, W. D. Jr. (1998). Water masses and mixing near the Antarctic Slope Front. In: S. S. Jacobs, & R. F. Weiss (Eds). Ocean, Ice and Atmosphere. Interactions at the Antarctic Continental Margin. Antarctic Research Series. Vol. 75, pp. 1–27. Whitworth, T., & Peterson, R. G. (1985). Volume transport of the Antarctic Circumpolar Current from bottom pressure measurements. J. Phys. Oceanogr., 15, 810–816. Wong, A. P. S., Bindoff, N. L., & Church, J. (1999). Large-scale freshening of intermediate waters in the Paciﬁc and Indian oceans. Nature, 400, 440–443, doi: 10.1038/22733. Wunsch, C. (2002). What is thermohaline circulation? Science, 298, 1179–1181. Wyrtki, K. (1961). The thermohaline circulation in relation to the general circulation in the oceans. Deep Sea Res., 8, 39–64.

Developments in Earth & Environmental Sciences, 8 F. Florindo and M. Siegert (Editors) r 2009 Elsevier B.V. All rights reserved DOI 10.1016/S1571-9197(08)00005-0

Chapter 5

Cenozoic Climate History from Seismic Reﬂection and Drilling Studies on the Antarctic Continental Margin Alan K. Cooper1,, Giuliano Brancolini2, Carlota Escutia3, Yngve Kristoffersen4, Rob Larter5, German Leitchenkov6, Phillip O’Brien7 and Wilfried Jokat8 1

Department of Geological and Environmental Sciences, Stanford University, Stanford, CA 94306, USA 2 Istituto Nazionale di Oceanograﬁa e di Geoﬁsica Sperimentale, B.go Grotta Gigante 42/c, 34010 Sgonico, Trieste, Italy 3 Instituto Andaluz de Ciencias de la Tierra, CSIC-Univ. de Granada Campus de Fuentenueva s/n, 18002 Granada, Spain 4 Department of Earth Science, University of Bergen, Alle`gaten 41, N-5007 Bergen, Norway 5 British Antarctic Survey, High Cross, Madingley Road, Cambridge CB3 0ET, UK 6 Research Institute for Geology and Mineral Resources of the World Ocean, VNIIOkeangeologia, 1 Angliysky Avenue, 190 121 St. Petersburg, Russia 7 Geoscience Australia, Cnr Jerrabomberra Avenue and Hindmarsh Drive, GPO Box 378, Canberra ACT 2601, Australia 8 Alfred-Wegener Institute, Am Handelshafen 12, D-27570 Bremerhaven, Germany

ABSTRACT Seismic stratigraphic studies and scientiﬁc drilling of the Antarctic continental margin have yielded clues to the evolution of Cenozoic climates, depositional paleoenvironments and paleoceanographic conditions. This paper draws on studies of the former Antarctic Offshore Stratigraphy Project and others to review the geomorphic and lithostratigraphic offshore features that give insights into the long-duration (m.y.) and short-term (k.y.) changes that document the great variability of Cenozoic Antarctic paleoenvironments. The lithologic drilling record documents non-glacial (pre-early Eocene) to full-glacial (late Pliocene to Holocene) times, and documents times of cyclic ice-sheet ﬂuctuations at k.y. Corresponding author. Tel.: þ01 650329 5157

E-mail: [email protected] (A.K. Cooper).

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scales (early Miocene to Pliocene and Holocene). Times of signiﬁcant change in types and/or amounts of glaciation are also seen in the offshore lithologic record (early Oligocene, mid-Miocene, early Pliocene). Seismic data illustrate large-scale geomorphic features that point to massive sediment erosion and dispersal by ice sheets and paleoceanographic processes (e.g. cross-shelf troughs, slope-fans, risedrifts). The commonality of these features to East and West Antarctica since late Eocene time points to a continent that has been intermittently covered, partially to completely, by glaciers and ice sheets. The greatest advances in our understanding of paleoenvironments and the processes that control them have been achieved from scientiﬁc drilling, and future progress depends on a continuation of such drilling.

Citation to regional-section information: Regional-section authors, 2008, In Cooper et al., Cenozoic climate history from seismic-reﬂection and drilling studies on the Antarctic continental margin, In: F. Florindo and M. Siegert (Eds). Antarctic Climate Evolution, Developments in Earth and Environmental Sciences, Vol. 8, Elsevier, 537p. Citation to general summary: Cooper, A.K., G. Brancolini, C. Escutia, Y. Kristoffersen, R. Larter, G. Leitchenkov, P. O’Brien, W. Jokat, 2008, Cenozoic climate history from seismic-reﬂection and drilling studies on the Antarctic continental margin, In: F. Florindo and M. Siegert (Eds). Antarctic Climate Evolution, Developments in Earth and Environmental Sciences, Vol. 8, Elsevier, 537p.

5.1. Introduction The Antarctic continental margin is a tectonic collage of former rifts and subduction zones that are covered by sediments deposited when the adjacent continent was free of regional glaciers (i.e. ‘pre-ice-sheet’ times) and when glaciers extended onto the margin (i.e. glacial times). Since the 1960s, many seismic surveys and sea-ﬂoor cores and a few drill cores have been acquired on the continental margin to decipher the Cenozoic and earlier history of Antarctica’s paleoenvironments and paleoclimates – a history hidden onshore in sediments now unreachable beneath the Antarctic Ice Sheet. This chapter summarizes principally key results of seismic and drilling studies for proximal parts of the continental margin done from 1989 to 2004 by the multinational Antarctic Offshore Statigraphy project (ANTOSTRAT) to decipher Antarctic Ice Sheet history. We include some ﬁndings of the successor Antarctic Climate Evolution project (ACE) that includes the Cenozoic Antarctic Stratigraphy and Paleobathymetry project (CASP), to create a uniﬁed circum-Antarctic stratigraphy from all existing seismic and rock-core data (Davey and Cooper,

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2007). Our summary complements istotopic and ice-rafting studies for distal parts of the margin and abyssal areas (e.g. Warnke et al., 1996; Zachos et al., 2001). We ﬁrst describe work in ﬁve geographic sectors of the margin, and then summarize key results for the entire margin. Multichannel seismic (MCS) reﬂection data, the principal tool for deep stratigraphic studies of the continental margin, have been recorded by more than 15 nations (Fig. I-1). Many topical MCS studies with maps of select regional data exist (see citations in regional sections below), but few comprehensive data compilations are either published or openly accessible. Notable exceptions are drilling results (e.g. Deep Sea Drilling Project, Ocean Drilling Program, Cape Roberts Project, ANDRILL and other drilling projects), MCS data compilations in the Antarctic Seismic Data Library System (e.g. www.scar-sdls.org; Wardell et al., 2007), a few Antarctic and regional geosciences atlases (e.g. Hayes, 1991; Cooper et al., 1995), online

Figure I-1: Map showing locations of tracklines for multichannel seismicreﬂection data on the Antarctic continental margin as of late 2006 (from Wardell et al., 2007). Regions are RS, Ross Sea; WL, Wilkes Land; PB, Prydz Bay; WS, Weddell Sea; AP, Antarctic Peninsula.

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data centres (e.g. World Data Center) and select discussions of Antarctic margin databases (e.g. Anderson, 1999). In general, Antarctica had a relatively warm climate and normal-waterdepth continental shelf (i.e. like low-latitude continental shelves today) in earliest Cenozoic and Late Cretaceous times – conditions that differ from the polar climate of the latest Cenozoic, with its thick ice sheet and an abnormally deep-water-depth and landward sloping shelf (e.g. Cooper et al., 1991b; Anderson, 1999). Ice has played an important role in continental margin evolution by eroding onshore areas (formerly with vegetation) and discharging the debris into the sea, where ocean currents distribute it to the continental shelf, slope and rise. At times the ice has strongly eroded parts of the shelf. Tectonic processes, principally variable thermal and ﬂexural subsidence and uplift, have also modiﬁed the margin morphology and hence stratigraphy (e.g. ten Brink et al., 1995). Offshore Antarctic stratigraphic studies have thus focused on mapping geomorphology and seismic facies of characteristic features (e.g. shelf-edge fans/deltas, mound deposits, unconformities, etc.), and using the limited core and downhole data to decipher their depositional paleoenvironments and relation to nearby ice, if any. Such features help to infer and establish where and when non-glacial and glacial processes acted (e.g. Cooper et al., 1991b). Drilling is the only way to ‘ground truth’ the regional seismic surveys (i.e. via a direct tie of lithologic facies to seismic facies), and to provide the age and biostratigraphic control needed to decipher depositional and climatic paleoenvironments (e.g. Barker and Camerlenghi, 2002; Cooper and O’Brien, 2004). The following regional subchapters have been written by the regional experts listed. Their bibliographic citations are augmented in a ‘selected reference’ section that provides additional background on the prior studies done by the Antarctic geoscience community.

5.2. Ross Sea (G. Brancolini and G. Leitchenkov) The Ross Sea has four large sedimentary basins with thick Cenozoic sequences that record the proximal paleoenvironmental histories of the East and West Antarctic Ice Sheets (Cooper and Davey, 1985; Cooper et al., 1991b, c). Here, ice-sheet evolution is linked to the Cenozoic uplift histories of the Transantarctic Mountains and Marie Byrd Land. Offshore, the West Antarctic Ice Sheet (WAIS) ﬂows across the Eastern basin, and the East Antarctic Ice Sheet (EAIS) passes over the Transantarctic Mountains and ﬂows across the Victoria Land basin, the Northern basin and the Central trough (Fig. RS-1).

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Figure RS-1: Map showing the structural framework of the ross sea. The limits of the basins are based on the depositional limits of the seismic Sequence RSS-1. VL, Victoria Land; VLB, Victoria Land basin; NB, Northern basin; CoH, Coulman high; CH, Central high; CT, Central trough; EB, Eastern basin; WAIS, West Antarctic Ice Sheet; EAIS, East Antarctic Ice Sheet. The heavy dashed line marks the postulated boundary between East Antarctic Ice Sheet and West Antarctic Ice Sheet drainage. Numerous seismic studies have been done in the Ross Sea region since the 1960s, with more than 45,000 km of MCS reﬂection data collected since 1980 (Hinz and Block, 1984; Sato et al., 1984; Cooper and Davey, 1987; Hinz and Kristoffersen, 1987; Zayatz et al., 1990; Brancolini et al., 1991) (Fig. RS-2a), to provide tectonic and deep stratigraphic control. A large number of singlechannel seismic (SCS) surveys have also been conducted for greater resolution of the shallow subsurface (Fig. RS-2b). Drilling at several sites by DSDP (Deep Sea Drilling Project, Hayes and Frakes, 1975), DVDP (Dry Valley Drilling Project, McGinnis, 1981), MSSTS (McMurdo Sound Sediment and Tectonic Study, Barrett, 1986), CIROS (Cenozoic Investigation in the Western Ross Sea, Barrett, 1989) and CRP (Cape Roberts Project,

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Figure RS-2: (A): Multichannel seismic-reﬂection surveys in the Ross Sea (modiﬁed from Brancolini et al., 1995). Some of these data are available in digital format from Cooper et al. (1995) and others from the Antarctic Seismic Data Library (Childs et al., 1994; Wardell et al., 2007). (B) Singlechannel seismic surveys in the Ross Sea (modiﬁed from Barrett et al., 1999).

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Figure RS-3: Correlation of Ross Sea area drilling and seismic stratigraphy with global oxygen isotope (Miller et al., 1987) and eustacy (Haq et al., 1987) curves (modiﬁed from Brancolini et al., 1995; Cooper et al., 1995). Regional erosional unconformities in Oligocene and younger sections are interpreted to be due in part to sub-ice erosion, especially in late Neogene time.

Cape Roberts Science Team, 1998, 1999, 2000, 2001) ANDRILL (Antarctic geological Drilling, Naish et al., 2007; Florindo et al., 2008; Harwood et al., 2008) provides geologic ground truth data (Fig. RS-3). A regional seismic stratigraphy has been derived by the ANTOSTRAT project with seismic sequences and unconformities tied to drilling data (Fig. RS-3; Cooper et al., 1995). Ross Sea seismic data are used by many to infer glacial sedimentary processes (e.g. Cooper et al., 1991b; Alonso et al., 1992; Anderson and

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Bartek, 1992; Shipp et al., 1994; Brancolini et al., 1995; Cochrane et al., 1995; De Santis et al., 1995; Bartek et al., 1996; De Santis et al., 1999; Bart et al., 2000; Bart, 2003; Chow and Bart, 2003; Accaino et al., 2005). Characteristic features and inferred processes in Oligocene and younger strata include: 1. Landward-deepening seaﬂoor of the continental shelf with broad (up to 100 km wide) cross-shelf troughs and banks formed by ice-stream erosion and deposition, respectively. 2. Numerous regional seismic unconformities believed to result from erosion of the continental shelf by grounded ice sheets. 3. Steep prograding sedimentary sequences (i.e. foreseest dips more than 51 and eroded topset strata) interpreted as ice-proximal till deltas from grounded ice. 4. Wedge-shape, non-reﬂective units interpreted as ‘till tongues’ deposited by grounded ice. 5. Shallow sediment with high velocities, considered due to overcompaction by grounded ice. The seismic stratigraphy and drilling help establish ice-sheet evolution in the Ross Sea region, and are discussed below for four key intervals.

5.2.1. Pre-Ice-Sheet (Pre-Late-Oligocene Time) This period includes seismic sequence RSS-1 between the acoustic basement and unconformity RSU-6 (Fig. RS-3, Foldouts RS-1 and RS-2) and, in the Victoria Land basin, acoustic units V4 and V5 (Cooper et al., 1987). Acoustic basement rocks have been sampled at two sites, and are Palaeozoic Beacon Formation rocks at the CRP-3 site adjacent to the coast (Cape Roberts Science Team, 2000) and are inferred palaeozoic and Cretaceous igneous and metamorphic rocks at DSDP Site 270 in the centre of the Ross Sea (Hayes and Frakes, 1975). The basins are believed to hold sedimentary rocks of Cretaceous and younger age (Hinz and Block, 1984; Cooper et al., 1991c), but Sedimentary rocks older than late Eocene have not been cored by drilling. Upper Eocene sediments have been cored in the CIROS-1 drillhole in McMurdo Sound (Coccioni and Galeotti, 1997; Fielding et al., 1997; Hannah et al., 1997; Monechi and Reale, 1997). The presence of ubiquitous lonestones (Barrett, 1989) testiﬁes that glaciers (but not necessarily continent-size ice sheets) were calving at sea-level then. Eocene erratic rocks are found in coastal areas (Levy et al., 1995), and have ﬂora indicating cool, but not glacial, climates in the McMurdo area (Stillwell and Feldmann, 2000). Offshore basin analysis, mainly from MCS

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reﬂection data, suggests that in Cenozoic pre-ice-sheet times, the Ross Sea was dissected by high-standing subaerial ridges, now seen as the buried Coulman high and Central high (De Santis et al., 1995). Prior outcropping of these ridges is suggested by the presence of regoliths above the basement at DSDP Site 270 (Hayes and Frakes, 1975). 5.2.2. Early Glacial (Late Oligocene to Early Miocene) This period includes seismic sequences RSS-2 and -3 (Fig. RS-3 and Foldouts RS-1 and RS-2). Sedimentary rocks from this period were recovered at CIROS-1, and CRP-1, -2 and -3, and MSSTS-1 drilling sites in the McMurdo Sound area. Such rocks include compacted diamicton indicative of deposition by/under grounded ice, as well as mud and ice-rafted debris (IRD) indicative of open-water environments (Barrett, 1986; Barrett, 1989; Hannah, 1994; Cape Roberts Science Team, 2001), in lower Miocene sediments at CRP 2/2A sites. Compacted diamicton and mud layers at site CRP-1, vary with uniform cyclicity, and document systematic oscillation of the EAIS size (Naish et al., 2001). The oscillations are at orbital periodicities similar to those recorded by isotope studies in distal deep-ocean sediments. Seismic facies along the border of the Victoria Land basin suggest that tidewater glaciers all along the Transantarctic Mountains intermittently extended onto the continental shelf and carried abundant glacial sediment to the sea (Brancolini et al., 1995; Bartek et al., 1996; Henrys et al., 2001). In the eastern Ross Sea, at DSDP Site 270, Nothofagus-dominated ﬂora in lower Miocene sediments (Kemp and Barrett, 1975) are similar to those recovered in McMurdo Sound drillcores (Hill, 1989; Mildenhall, 1989; Askin and Raine, 2000), and indicate cool-temperate climates during interglacial periods. DSDP Site 270 also recovered lower Miocene iceproximal glaciomarine sediments from the early Miocene section, but the size and character of the ice sheet that deposited these sediments is debated. Anderson and Bartek (1992) suggest, based on high-resolution singlechannel seismic data and drill cores, that by late Oligocene to early Miocene time, the continental shelf was deeply scoured and foredeepened (i.e. landward dipping) by a massive ice sheet. In contrast, Brancolini et al. (1995) and De Santis et al. (1995), utilize regional stratigraphic maps (Cooper et al., 1995) and their seismic facies analyses to postulate that, during the same period the Central high was partly exposed and partly covered by small subpolar ice caps (i.e. subpolar as deﬁned by Anderson and Ashley, 1991). A semi-quantitative evaluation of the water depth of the Eastern basin during the early Miocene, based on the backstripping of the seismic section

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in Foldout RS-1, indicates that the foredeepeened proﬁle of the Eastern basin was only attained after middle Miocene time (De Santis et al., 1999). The end of this early glacial period is marked by a change in reﬂection geometries beneath the outer continental shelf (Foldouts RS-1 and RS-2) from principally aggrading (RSS-2 and -3) to principally prograding, (RSS-4). Cooper et al. (1991b) postulate that this change marks the start of grounded-glacier advances to the shelf edge and erosion of a normalwater-depth shelf by episodic grounded ice. 5.2.3. The Ice-Sheet Development (Mid-Early Miocene to Early Pliocene) This period includes seismic sequences RSS-4, -5 and -6 (Fig. RS-3 and Foldouts RS-1 and RS-2). Sediment from these sequences was recovered at DSDP Sites 271, 272 and 273, MSSTS-1 and DVDP 10/11. The early-middle Miocene is postulated to have been a time of major ice buildup of ice in the Ross Sea region, and the carving of the ﬁrst deep troughs, similar in size to the present ice streams, across the continental shelf (Anderson and Bartek, 1992). Bart (2003) and Chow and Bart (2003) recognize at least two major WAIS expansions during the early part of the middle Miocene and ﬁve in the entire Miocene. These expansions suggest that either portions of the West Antarctic land elevation were above sea-level and/or the air and water temperatures were sufﬁciently cold to support a marine-based ice sheet. Drill cores from the middle Miocene have been recovered at DSDP Sites 272 and 273 (Hayes and Frakes, 1975; Savage and Ciesielsky, 1983; Leckie and Webb, 1986), and consist of diatom-bearing sediments interpreted as waterlain tills and proximal- to distal-glacimarine deposits (Hambrey and Barrett, 1993). Upper Miocene rocks are missing from all continental shelf drill cores, except in the McMurdo Sound region (MSSTS-1, DVDP-10 and -11 drill sites), where glaciomarine diamictites (tillites) and terrestrial strata are found. These deposits are interpreted as having originated from glaciers ﬂowing out of the Transantarctic Mountains (Powell, 1981; Barrett, 1986; Ishman and Webb, 1988; McKelvey, 1991). The recent ANDRILL drilling on the Ross Ice shelf, near McMurdo Sound (MIS project; Naish et al., 2007), recovered a 1,284 m long core that records Antarctica’s history over the last 14 million years. The core indicates periods of ice-sheet growth, advancing over the drill site and then retreating again to allow the open-marine conditions to return. More than 60 of these advance-retreat cycles are present in the core. On the outer continental shelf and upper slope, well-stratiﬁed seismic sequences inferred to be of late Miocene age (i.e. RSS-6) are present in the

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Northern and Eastern basins, but the sequences are thin or absent on the inner shelf (Cooper et al., 1995). In both basins, these sequences are characterized by steeply prograding clinoforms with relatively thin or eroded topset beds and a major seaward shift of the palaeo-shelf edge (Foldouts RS-1 and RS-2). The sequences are thought to have been deposited by intermittent grounded ice sheets carrying sediment to the continental shelf edge (Bartek et al., 1991; Cooper et al., 1991b; Anderson and Bartek, 1992). 5.2.4. The Polar Ice Sheet (Early Pliocene Through Quaternary) This period includes seismic sequences RSS-7 and -8 (Fig. RS-3) and seismic units 1–7 of Alonso et al. (1992) and Anderson and Bartek (1992) (Fig. RS-4a and b). Pliocene sediment has been recovered at drill sites in the Taylor Valley (DVDP-10 and -11) and on the continental shelf at DSDP Sites 271 and 273. DVDP drill cores contain sediments deposited by the Taylor Glacier, while the sparsely sampled Pliocene deposit at DSDP 271 and 273 contains diatomaceous glaciomarine strata (Hayes and Frakes, 1975). These rocks imply that warmer interglacial conditions than today existed at that time (Anderson and Ashley, 1991). Seismic sequences of inferred Pliocene through Quaternary age occur in the EAIS drainage in the Northern basin, where they are up to 800 m thick in the till delta fan system (Cooper et al., 1995). In the Eastern basin, Pliocene through Quaternary age strata lie within the WAIS drainage and are more than 1,000 m thick (Cooper et al., 1995). Detailed seismic stratigraphic analyses from the Eastern basin margin (Fig. RS-4a; Alonso et al., 1992; Fig. RS-4b; Anderson and Bartek, 1992), recognize a major change in the seismic character of the Pliocene deposits. Up-section, the seismic unit thicknesses decrease, the geometry of the sequences changes from principally progradational to aggradational, and numerous widespread glacial erosion surfaces are seen. These features indicate more frequent grounding events on the continental shelf and increased subglacial till deposition relative to basal transport of sediments to the grounding line. Bart et al. (2000) and Anderson et al. (1992) suggest that on at least eight occasions during Pliocene to Quaternary times, the East and West Antarctic Ice Sheets were much larger than today. The frequent and extensive grounding events on the outer continental shelf contradict the widely held view that the land-based EAIS was relatively stable and the largely marine-based WAIS was relatively dynamic (Bart and Anderson, 2000). The last glacial maximum (LGM) in the Ross Sea has been studied using seaﬂoor cores, subbottom and swath bathymetry data (Thomas and Bentley, 1978; Kellogg et al., 1979; Denton et al., 1989; Leventer et al., 1993; Brambati et al., 1994; Hilﬁnger et al., 1995; Kellogg et al., 1996; Licht et al.,

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Figure RS-4: Seismic-reﬂection proﬁles across the Eastern basin (A) and the Northern basin (B), illustrating the Neogene stratigraphic sections (from Anderson and Bartek, 1992). The shelf margin delta fan complex shown is a common feature in the seismic data from the continental margin and characterizes deposition close to the ice-sheet grounding line. See Fig. RS-1 for location. 1996; Cunningham et al., 1999; Domack et al., 1999; Licht et al., 1999; Shipp et al., 1999; Alley and Bindschadler, 2001). Radiocarbon dates from diamictons and sediment composition, indicate that ice-free conditions existed on the inner shelf at times during the period from 60 to 10 ka, and that the ice sheet was present between 26.5 and 19.5 ka. (Domack et al., 1999). The maximum ice-sheet expansion in the LGM is still debated: Kellogg et al. (1996) place the grounding line at the continental shelf edge, Domack et al. (1999) put it just north of the Coulman Island and Licht et al. (1999) interpret that it was about 100 km south of Coulman Island. The last retreat of the grounding line occurred in the western Ross Sea around 11 ka at a rate of about 100 m/year, and the grounding line reached its present position about 6 ka (Domack et al., 1999; Shipp et al., 1999).

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5.3. Wilkes Land (C. Escutia and P. O’Brien) 5.3.1. Acoustic Stratigraphy The approximately 1,500 km long Wilkes Land segment of the continental margin (Fig. WL-1) formed during the Cretaceous separation of Australia and Antarctica (Cande and Mutter, 1982; Veevers, 1987; Sayers et al., 2001; Colwell et al., 2006; Leitchenkov et al., 2007; O’Brien and Stagg, 2007). The stratigraphy of the margin is known mainly from the seismic stratigraphic interpretation of numerous MCS surveys (Sato et al., 1984; Wanneson et al., 1985; Tsumuraya et al., 1985; Eittreim and Hampton, 1987; Ishihara et al., 1996; Tanahashi et al., 1997; Brancolini and Harris, 2000; Stagg et al., 2004a, b); complemented by surface sediment cores (Domack et al., 1980; Payne and Conolly, 1972; Domack, 1982; Tsumuraya et al., 1985; Hampton et al., 1987; Ishihara et al., 1996; Tanahashi et al., 1997; Brancolini and Harris, 2000; Leventer et al., 2001; Escutia et al., 2003; Michel et al., 2006); and limited deep geological sampling recovery at DSDP Sites 268 and 269 (Hayes and Frakes, 1975). The best-surveyed area is the eastern Wilkes Land margin (EWL) from the Ade´lie Coast to George V Land. West of this area (the western Wilkes Land margin-WWL), Japan and Russia collected widely spaced seismic lines that were then augmented during the 2001–2002 Australian Antarctic and Southern Ocean Proﬁling (ASSOPP) Project (Stagg et al., 2004a, b; Leitchenkov et al., 2007).

5.3.1.1. Pre-ice-sheet stratigraphy Along the Wilkes Land margin syn- and post-rift sedimentary rocks reach thicknesses of more than 7 km (Stagg et al., 2004a, b). Pre-Eocene syn-rift strata are about 3 km thick and are highly variable in seismic character, with discontinuous, faulted, and tilted strata onlapping the ﬂanks of the acoustic basement (Eittreim and Smith, 1987; Eittreim, 1994; De Santis et al., 2003; Stagg et al., 2004a, b; Leitchenkov et al., 2007). The thickest (at least 9 km) depocentre of post-rift sedimentary rocks is located in the WWL off the Bud Coast (Close et al., 2007). In the EWL postrift strata are up to 5 km thick across the Wilkes Land continental shelf, slope and rise (Eittreim and Smith, 1987; Hampton et al., 1987; Wannesson, 1990; Tanahashi et al., 1994; De Santis et al., 2003). These strata are welllayered on the continental rise and become less stratiﬁed and more discontinuous landward (Eittreim and Smith, 1987; Eittreim, 1994;

Surveys France (IFP) 1982 US Geological Survey 1984 Japan (JNOC) 1983-1996 Italian-Australian 2000 Australian 2000-2002 Russia (RAE) 2005-2007

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Figure WL-2: Multichannel seismic line WEGA W21 and line drawings of multichannel seismic proﬁles IFP 107 and WEGA W35 showing the overall architecture of the Wilkes Land margin from the continental shelf to the continental rise (modiﬁed from Escutia et al., 2005). See Fig. WL-1 for location of the seismic lines. De Santis et al., 2003). A prominent regional unconformity (WL-U3) within the Cenozoic post-rift section beneath the continental shelf (Fig. WL-2) is believed to be due to erosional processes related to the ﬁrst advance of grounded ice sheets onto the continental shelf (Eittreim and Smith, 1987; Tanahashi et al., 1994; Eittreim et al., 1995; Escutia et al., 1997; Escutia et al., 2005). The pre-ice-sheet strata below unconformity WL-U3, where resolvable, are ﬂat-lying and less stratiﬁed than glacial strata above the unconformity. Pre-ice-sheet rocks have been dredged from the Wilkes Land continental shelf and slope. On the inner shelf, Mesozoic sediments have been exposed via erosion by late Cenozoic glaciers near the Mertz ice tongue. Lignite was recovered (Mawson, 1940, 1942), and lower Cretaceous brecciated, carbonaceous siltstone was cored (Domack et al., 1980). Other dredge samples in the area, acquired by Leventer et al. (2001), include sedimentary clasts of Paleogene lignites with reworked Early Cretaceous ﬂora. On the upper continental slope off Terre Ade´lie, Sato et al. (1984) dredged samples

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of locally derived Oligocene and Miocene limestone and undated sedimentary, metamorphic and igneous rocks of mostly ice-rafted origin.

5.3.1.2. Continental shelf glacial stratigraphy Glacial sequences on the shelf thicken seaward in prograding wedges (Fig. WL-2). The sequences are deeply eroded by broad troughs that cross the shelf. The troughs are interpreted as the erosional paths of ice streams during times of glacial maxima (Eittreim et al., 1995). Foreset strata are commonly truncated at or near the seaﬂoor beneath the troughs (Fig. WL-2). Topset strata form the banks adjacent to the troughs. Ice is inferred to have moved slowly over bank areas and rapidly in the troughs. Geometry of strata in buried troughs on the shelf suggests to some (Eittreim et al., 1995; Escutia et al., 2000) that the locations of ice streams and their erosional troughs and banks have shifted during consecutive glacial advances. Regional glacial seismic sequences and unconformities deﬁned by different workers (Table WL-1) in the EWL were renamed and in some cases redeﬁned by De Santis et al. (2003). On the shelf, sequences are truncated by two Table WL-1: Summary of the terminology assigned in previous publications to the inferred wilkes Land glacial sequences and their bounding unconformities (updated from Escutia et al., 2005). Unconformities (tied with lines) and sequences (in between these lines) are listed from younger at the top to older.

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regional unconformities, WL-U3 and WL-U8 (Wannesson et al., 1985; Eittreim and Smith, 1987; Hampton et al., 1987; De Santis et al., 2003), and the erosion is thought to result from grounded ice sheets moving across the continental shelf (Tanahashi et al., 1994; Eittreim et al., 1995; Escutia et al., 1997; Escutia et al., 2005). Eittreim et al. (1995) calculated erosion of 300 to 600 m of strata below WL-U3. Sequences below WL-U8 are dominantly aggradational and sequences above are principally progradational. For unconformity WL-U8, Eittreim et al. (1995) estimated erosional truncation of 350 to 700 m of sediment. Unconformity WL-U8 marks changes in the geometry of the outer shelf progradating wedges, from shallower dips below WL-U8 to steeper dips above (foreset slopes up to about 101). During the open-marine Holocene, thick laminated diatom mud and oozes were deposited in deep (W1,000 m) inner shelf basins, such as for example the Ade´lie Drift (Costa et al., 2007). Based on AMS radiocarbon dates, this drift has accumulation rates on the order of 20–21 m/k.y. Opal, Ti and Ba timeseries show decadal to century variance suggestive of solar forcing and El Nin˜o Southern Oscillation (ENSO) forcing (Costa et al., 2007).

5.3.1.3. Continental slope glacial stratigraphy Although partly obscured by seaﬂoor multiples, the stratigraphy of the continental slope consists of seaward-dipping reﬂectors (Eittreim and Smith, 1987; Hampton et al., 1987; Eittreim et al., 1995). Prograding strata above the WL-U8 unconformity downlap and pinch out at the base of the continental slope, but deeper units (i.e. between WL-U8 and WL-U3) continue across the margin (Hampton et al., 1987; Eittreim et al., 1995; Escutia et al., 1997; De Santis et al., 2003) (Fig. WL-2). Sediments forming prograding foresets were delivered directly to the outer shelf and upper slope as deforming tills at the base of ice streams at times of glacial maxima (Eittreim et al., 1995). Ice-stream delivery of a large volume of unconsolidated sediment to the steep slope resulted in sediment failures that led to the development of large chaotic deposits at the base of the paleoslope foresets (Eittreim et al., 1995; Escutia et al., 2000; De Santis et al., 2003; Escutia et al., 2007). More-recent slope strata are dissected by erosional submarine gullies (Eittreim et al., 1995) and slope canyons (Escutia et al., 2000). Sea-ﬂoor sediment cores from the continental slope contain debris-ﬂow units and numerous hiatuses. The oldest sediment has been dated as late Miocene in age, indicating that gravity ﬂows have been a dominant slope process since at least this time (Escutia et al., 2003).

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5.3.1.4. Continental rise glacial stratigraphy On the EWL continental rise, strata above the WL-U3 unconformity include six glacial-related seismic units, WL-S4 to WL-S9 (De Santis et al., 2003; Donda et al., 2003) (Table WL-1, Fig. WL-3). The two deepest units, WL-S4 and WL-S5, consist of stratiﬁed and continuous reﬂectors that onlap at the base of the slope (Escutia et al., 1997; Donda et al., 2003). Acoustic signatures of isolated channel-levee complexes that characterize turbidite deposition are ﬁrst observed up-section within unit WL-S5 (Escutia et al., 1997; Escutia et al., 2000; Escutia et al., 2002; Donda et al., 2003). During deposition of units WL-S6 and WL-S7, channel-levee complexes became widespread and turbidity ﬂows were the dominant process building the sedimentary ridges on the rise. Wavy reﬂectors that are characteristic of bottom contour-current deposition occur on the lower rise in unit WL-S6 and on the upper rise in WL-S7. Within Unit WL-S8, there is evidence for bottom contour-current and turbidite ﬂows, but WL-S8 mostly inﬁlls previous depressions (Escutia et al., 1997; Escutia et al., 2000; Escutia et al., 2002; Donda et al., 2003). Unit WL-S9 is a discontinuous unit on the rise, and, where present, comprises channel and levee complexes and layered reﬂectors (Donda et al., 2003). Recent studies on the WWL margin glacial strata show a similar evolution of the glacial sedimentary sequences (i.e. increased proximal turbidite facies up-section and inﬂuence of bottomcontour-current deposition) above unconformities ‘eoc’ (Close et al., 2007) and WL3 (Leitchenkov et al., 2007), which correlate with WL-U3 on the 4.0

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EWL. Between 1101 and 1301 large debris-ﬂow deposits are also reported forming throughout the Miocene (Donda et al., 2007a, b). 5.3.2. Drilling on the Wilkes Land Margin DSDP Leg 28 drilled Sites 268 and 269 on the continental rise and abyssal plain, respectively, to determine the geologic and climate history of Antarctica and the Southern Ocean (Hayes and Frakes, 1975). The drill cores document that extensive Antarctic glaciation began at least by Oligocene to early Miocene time, and that water temperatures were cooltemperate in the late Oligocene and early Miocene and cooled during the Neogene, presumably as glaciation intensiﬁed. DSDP Site 268 was drilled to a subbottom depth of 474.5 m in 3,544 m water depth with total core recovery of 14% (Hayes and Frakes, 1975). Three units were described, based on lithologies and amounts of diatoms, nannofossil ooze and ice-rafted pebbles and granules (Hayes and Frakes, 1975). Piper and Brisco (1975) interpreted the two deeper units to be contourites, based on the character of silt laminae. The shallowest of the three units, dated as Pliocene and Quaternary, was interpreted as turbidites, based on a high content of silty clay with common silt laminae and ﬁne-sand beds 2–20 cm thick. Hayes and Frakes (1975) infer that the deepest lower Miocene and Oligocene unit was deposited when the ice sheet ﬁrst advanced onto the shelf. Water at that time was warm enough to support calcareous biogenic sedimentation, but ice-rafting and contourites provide evidence for nearby ice on East Antarctica and for bottom currents, possibly generated by cold bottom water production associated with a limited ice shelf or tongue (Hayes and Frakes, 1975). DSDP Site 269 was drilled to a subbottom depth of 958 m in a water depth of 4,285 m and with 42% recovery of Eocene to recent rocks (Hayes and Frakes, 1975). The section consists predominantly of silts and clays with variable amounts of microfossils. Diatom oozes and diatom mud dominate the upper half of the section, which is dated as Quaternary to Late Miocene in age (Hayes and Frakes, 1975). In the lower half, which is late Miocene to early Miocene and Oligocene in age, diatoms are absent but calcareous nannofossils are found in trace amounts. Similar to DSDP Site 268, there is a transition in facies at DSDP Site 269 from more distal facies in the lower part of the core to more proximal facies near the surface. Piper and Brisco (1975) interpret this facies change as resulting from substantial increased supply of sand and coarse silt and clay from the Antarctic continent, possibly in response to prograding of the continental margin.

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5.3.3. The Inferred Long-Term Record of Glaciations Investigators interpret the WL-U3 unconformity as having been eroded during the ﬁrst grounding of an ice sheet on the continental shelf (Tanahashi et al., 1994; Eittreim et al., 1995; Escutia et al., 1997; Escutia et al., 2005), either about 40 m.y. ago (Eittreim et al., 1995) or 33.4–30 Ma (Escutia et al., 2005) (Table WL-2). Above WL-U3, early glacial strata (e.g. likely glacial outwash deposits) were provided by ﬂuctuating temperate glaciers, and were deposited as low-dip-angle prograding foresets. The increase in stratal dips across unconformity WL-U8 in the prograding wedge at the shelf edge is interpreted to record a glacier-regime change from intermittent ﬂuctuating glaciers to persistent oscillatory ice sheets, either on the Late Miocene (Escutia et al., 2005) or about 3 Ma (Rebesco et al., 2006) (Table WL-2). The steep foresets above WL-U8 likely consist of ice proximal (i.e. waterlain till and debris ﬂows) and open-water sediments deposited as grounded ice sheets Table WL-2: Continental shelf and rise stratigraphy and inferred East Antarctic Ice Sheet evolution in the Wilkes Land margin and timing of events.

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extended intermittently onto the outer shelf – similar to sediments recovered from ODP Site 1167 on the Prydz Trough fan (O’Brien et al., 2001). On the continental rise, the up-section increase in the energy of the depositional environment in units WL-S5 to WL-S7 (i.e. seismic facies indicative of proximal turbidites and of bottom-contour-current deposition) likely resulted from enhanced shelf progradation. Maximum rates of sediment delivery to the rise appear to have occurred during the development of units WL-S6 and WL-S7, which is inferred to have been during the Miocene (Hayes and Frakes, 1975; De Santis et al., 2003; Escutia et al., 2005). During deposition of WL-S8 and WL-S9, sediment supply to the lower continental rise decreased and depocentres shifted landward to the base of the slope and outer shelf (Escutia et al., 2002; De Santis et al., 2003; Donda et al., 2003; Escutia et al., 2005; Rebesco et al., 2006). Inferred age for Units WL-S8 and WL-S9 is Pliocene to Recent (De Santis et al., 2003). Sequence WL-S9 was deposited under a polar regime with a persistent ice sheet during the Pliocene–Pleistocene. At that time, most sediment delivered to the margin was trapped on the outer shelf and slope, forming steep prograding wedges, with some sediment bypassing the slope in channelized turbidity currents (Escutia et al., 2002; De Santis et al., 2003; Escutia et al., 2005). During the Holocene open-water interglacial thick sections of diatom mud and oozes are deposited in deep inner shelf basins (Costa et al., 2007). These sediments hold an ultra-high-resolution record of climate variability likely by solar and ENSO forcing.

5.4. Prydz Bay (P. O’Brien and G. Leitchenkov) Prydz Bay is a re-entrant in the East Antarctic margin, and overlies a rift structure that extends about 500 km into the interior of the continent. The rift has channelled drainage at least since the Early Cretaceous (Fig. PB-1; Arne, 1994) and presently controls the Amery Ice Shelf drainage system, which drains more than 16% of East Antarctica. This drainage basin includes the Gamburtsev Mountains, a subglacial range in which the Cenozoic ice sheet may have nucleated. Its long history, thick sediment and Cenozoic outcrops in the ﬂanking Prince Charles Mountains have made Prydz Bay a likely site for preserving palaeo-climate records. Seismic surveys by Australia, Russia, Japan and the US and two ODP Legs 119 and 188 (Fig. PB-1), plus ﬁeld studies and exposure dating, have provided an extensive picture of palaeo-climate evolution of the region.

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5.4.1. Pre-Ice-Sheet (Pre-Late Eocene) The Lambert Graben and Prydz Bay basin formed during the Carboniferous or Permian (Arne, 1994; Lisker et al., 2005) and were depocentres in pre-icesheet times. Seismic data from the Prydz Bay shelf show pre-ice-sheet sequences of parallel, moderately continuous reﬂectors (Figs. PB-2 and PB-3). ODP Sites 740, 741 and 1166 penetrated pre-ice-sheet sediments that were deposited in ﬂuvial to ﬂuvio-deltaic environments. ODP Site 740 intersected interbedded sandstone, siltstone and mudstone with reddish coloration (Shipboard Scientiﬁc Party, 1989), interpreted as ﬂuvial ﬂood plain deposits (Turner, 1991). The red coloration suggests a seasonal ﬂuctuating rainfall regime but the age of this unit remains unknown (Truswell, 1991). It could be as old as Triassic, based on the presence of Triassic sediments in the northern Prince Charles Mountains (Leitchenkov, 1991; McLoughlin and Drinnan, 1997a, b). However Leitchenkov (1991) identiﬁed a thick (up to 5 km), faulted and high-velocity (up to 5.2 km/s) unit on multichannel data underlying these red beds. This sequence predates the main phase of breakup-related crustal extension, leading him to correlate the deep unit with Permian-Triassic sediments of the northern Prince Charles Mountains (Leitchenkov, 1991). If so, then the red beds in ODP Site 740 are likely Early Cretaceous or Late Jurassic in age. ODP Sites 741 and 1166 intersected Cretaceous sediments beneath the Cenozoic section. The Cretaceous comprises interbedded dark siltstone and

Figure PB-2: Diagramatic section across Prydz Bay shelf and slope based on Russian seismic lines SAE 32002 and RAE 4005. K1 are Cretaceous sediments. Location shown in Fig. PB-1.

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Figure PB-3: Seismic section between ODP sites 742 and 1166. Sequence PS2A2 comprises ﬂuvio-deltaic sands of late Eocene age. Sequence PS2A1 overlies an erosion surface and comprises late Eocene marine muds with lonestones. Location shown in Fig. PB-1 (modiﬁed from Erohina et al., 2004). sandstone with minor coal of probable delta plain to lagoonal origin. The ODP Site 1166 section is Turonian-Santonian(?), whereas ODP Site 741 recovered an older section of middle Aptian sediment (Fig. PB-3). Macphail and Truswell (2004) describe the palynomorphs from ODP Site 1166, and interpret the assemblage as indicating a conifer-dominated woodland vegetation, consistent with a cool, humid climate. The continental rise seaward of Prydz Bay contains up to 5 km of post-rift sediments (Figs. PB-2, Foldouts PB-1 and PB-2; Mizukoshi et al., 1986, Stagg et al., 2004a, b). The lowermost seismic stratigraphic unit has parallel, mostly continuous reﬂectors typical of deep-ocean deposition that probably occurred during pre-ice-sheet times (Mizukoshi et al., 1986; Kuvaas and Leitchenkov, 1992; Kuvaas et al., 2005).

5.4.2. Early Glacial (Late Eocene) In Prydz Bay, ODP Sites 739, 742 and 1166 recovered sediments deposited immediately before major glaciation (Barron et al., 1991; Cooper and O’Brien, 2004). The lithologies vary from dark siltstones to poorly sorted sands and bedded mudstone with lonestones. Seismic sections show that the

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sands overlie an undulating erosion surface, suggesting a period of erosion, possibly related to a relative low stand of sea-level (Erohina et al., 2004, Fig. PB-3). The sand unit ﬁnes up-section into the mudstone, which contains lonestones, marine diatoms and dynocysts (Shipboard Scientiﬁc Party, 2001a). Strand et al. (2003) interpret the sand unit as a ﬂuvial to delta plain channel deposit. They found sand-grain surface textures that suggest erosion and breakage by glaciers, implying the presence of at least valley glaciers in the hinterland of Prydz Bay. The overlying mudstone with lonestones suggests a marine transgression, with ﬂoating ice as a feature of the resulting shallow embayment. Macphail and Truswell (2004) report palynomorphs in the ﬁne-grained units that indicate a late Eocene age (middle Nothofagites asperus zone), representing an age range from 33.9 to 39.1 Ma. This age overlaps with the age suggested by diatoms in the transgressive mudstones (33–37 Ma, Shipboard Scientiﬁc Party, 2001a). Macphail and Truswell (2004) also propose that the palynological assemblage was derived from a ﬂora similar to stunted Nothofagus rainforest scrub, and consisted of ground-hugging plants and canopy trees about 1 m high. Today, such ﬂoras occur outside of Antarctica at higher altitudes, where cool temperatures limit tree growth. Therefore, the Prydz Bay ﬂora reﬂects a cool to cold environment at sealevel. More precise temperature estimates are not possible because the plants present were tolerant of a wide range of conditions (Macphail and Truswell, 2004). 5.4.3. Ice-Sheet Development (Oligocene–Miocene) The older glacial section of the shelf comprises tabular units that pinch out shoreward due to inner-shelf erosion, and that extend seaward into prograding slope deposits (Cooper et al., 1991a). Palaeo-shelf edges for these units are better deﬁned up-section as foreset strata steepen seaward (Fig. PB-2, Foldout PB-1). Shelf drilling (ODP Sites 739, 740, 741, 742 and 1166) recovered probable subglacial and glacimarine diamicts, with thin interbedded diatomaceous mudstones deposited during warm episodes (Hambrey et al., 1991; Erohina et al., 2004). The drilling and seismic evidence indicates glacial advance well across the Prydz Bay shelf during cold episodes, probably reaching the shelf edge. Over-compacted horizons indicate periods of glacial erosion and ice loading during the early Oligocene, Miocene and Plio-Pleistocene (Solheim et al., 1991; Shipboard Scientiﬁc Party, 2001a). Before the late Miocene, the Prydz Bay shelf prograded uniformly across its width, with the bulk of the ice and entrained sediment

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coming from the southern end of the bay (i.e. from the Lambert Graben). The Prydz Bay continental slope became progressively steeper from the early phase of glaciation in early Oligocene time, to reach angles of as much as 81 on the present slope (Foldouts PB-1 and PB-2). On the continental rise, a pre-ice-sheet unit is overlain by one exhibiting channel-levee geometries. The nature of the change in geometry and the tracing of reﬂectors to the shelf drilling suggest that this change originated from the glacial expansion and increased sediment supply in the early Oligocene (Kuvaas and Leitchenkov, 1992). Overlying the channel-levee complexes are sequences that include thick mounds and sediment waves suggestive of contourite deposition, in addition to turbidite channels and associated levees formed by intensiﬁed down-slope and along-slope currents in the early Miocene (Fig. PB-2, Foldout PB-2). ODP Site 1165 (Leg 188) drilled 999 m with 69% recovery into a thick mound of lower Miocene and younger contourite sediments with turbidites only in the upper 5 m (Cooper and O’Brien, 2004). The hole penetrated the base of the mounded sequences, which was still of early Miocene age (Handwerger et al., 2004). The drilling conﬁrmed the seismic interpretation that deposition of the thick contourite mounds had commenced by at least early Miocene time, but sediments above and below the surface were typical contourites – ﬁssile claystones with abundant silt laminae (Handwerger et al., 2004). Therefore, there was no obvious lithological change in the hole to suggest a reason for the change from low relief submarine fans to highly mounded deposits, previously inferred to be mixed turbidite-contourites. ODP Site 1165 intersected a surface that can be mapped along the rise, and that marks a middle Miocene (14–16 Ma) change in sedimentation from laminated contourites to hemipelagic and pelagic facies (Cooper and O’Brien, 2004). Also, minerals and fossils recycled from shelf deposits ﬁrst appear, suggesting the start of intense erosion by ice and overdeepening of the shelf. At this time, sedimentation rates slow more rapidly at the drill site, falling from 100 m/m.y. in the early Miocene to 37 m/m.y. in the midMiocene to 10 m/m.y. during the Plio-Pleistocene (Shipboard Scientiﬁc Party, 2001c; Florindo et al., 2003; Fig. PB-4). On shorter time scales, Gru¨tzner et al. (2003) examine the proportions of terrigenous sediment and biogenic opal in ODP Site 1165 between 3.4 and 7.6 Ma. They ﬁnd high opal content from 5.8 to 5.2 Ma, which they relate to reduced sea ice and increased productivity. They also identify terrigenous intervals with high sedimentation rates from 7.2–6.6 Ma and 5.2–4.8 Ma, which they interpret as indicating high erosion rates and a ﬂuctuating ice sheet under the inﬂuence of obliquity forcing. Gru¨tzner et al. (2003) also report cyclic variations in sediment composition and physical properties that have

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Figure PB-4: Age-depth model for ODP Site 1165 from Shipboard Scientiﬁc Party (2001c) showing rapid sedimentation during the early Miocene, reducing rapidly through the late Miocene to Pliocene.

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spectral peaks at B94, 41, 31, 21, and 18 k.y. cycles. Williams and Handwerger (2005) report that geophysical log parameters detect cycles of biogenic and terrigenous input at periods of B15–23 and B135 k.y., probably representing Milankovich-scale forcing of paleoenvironmental processes. Uncertainty in the age model of the hole prevents them from exact matching of peaks. 5.4.4. The Polar Ice Sheet (Late Miocene(?)–Pleistocene) In the early Pliocene, ice ﬂow regimes changed and ice was focused into an ice stream on the western side of the bay, cutting a cross-shelf trough, the Prydz Channel. The ice stream delivered basal debris to the shelf edge, where the debris built a trough mouth fan on the upper continental slope (Prydz Channel Fan, O’Brien and Harris, 1996; O’Brien and Leitchenkov, 1997; O’Brien et al., 2004). On the banks adjacent to Prydz Channel, vertical aggradation of subglacial debris produced tabular units while glacial erosion overdeepened the inner shelf. Two ODP holes were drilled into the continental slope. ODP Site 743 was drilled to 98 mbsf into the eastern, steep part of the slope, and recovered diamict. ODP Site 1167 was drilled to 447.5 mbsf into the Prydz Fan, and recovered muddy, pebbly sands and diamicts deposited by slumping of subglacial debris interpreted to have originated at the ice grounding line at the shelf edge (Foldout PB-2, O’Brien et al., 2001; Passchier et al., 2003). ODP Site 1167 also recovered thin mudstone units deposited during periods of reduced ice extent (Shipboard Scientiﬁc Party, 2001b, Passchier et al., 2003). More than 90% of the fan was deposited before the mid Pleistocene, and there were only three advances of the Amery Ice Shelf to the shelf edge in the late Pleistocene (O’Brien et al., 2004). Clay mineralogy, magnetic properties and clast composition at ODP Site 1167 show changes suggesting that the Pleistocene peak of erosion and ice volume in the Lambert-Amery drainage system occurred in the early Pleistocene (O’Brien et al., 2004). Oxygen isotope measurements on foraminifera from ODP Site 1167 also suggest that sedimentation was reduced after the mid-Pleistocene, with the last ice advance to the shelf edge at about Marine Isotope Stage 16 (612–698 ka; Lisiecki and Raymo, 2005). However, the stratal record is fragmentary because hiatuses are common, which leads to a tentative identiﬁcation of isotope stages (Theissen et al., 2003). During the mid-to-late Pleistocene, ice advances were less extensive. During the last glacial cycle, the Amery Ice Shelf grounded only 100 km north of the present ice shelf edge and far from the continental shelf edge (Domack et al., 1998; O’Brien et al., 1999) (Fig. PB-1).

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On the continental rise, sedimentation rates decreased through the Pliocene and Pleistocene because less detritus was eroded from the continent and because sediment was deposited on the upper slope in front of the Prydz Channel. The inferred early Pliocene base of the Prydz Channel Fan is the prominent unconformity mapped by Mizukoshi et al. (1986, Reﬂector A) and O’Brien et al. (2004, Reﬂector PP-12). Drilling and seismic evidence indicate that glaciers advanced to the edge of the Prydz Bay shelf in cold episodes during the Pliocene and early Pleistocene, yet evidence of warm episodes also exists. Sediments in the Prince Charles Mountains indicate open-water fjordal environments in the Miocene to Pliocene (Hambrey and McKelvey, 2000; Whitehead et al., 2003, 2004). Lower Pliocene marine diatomite in the Vestfold Hills, on the eastern side of Prydz Bay, contains evidence of temperatures 41C warmer than today (Whitehead et al., 2001). ODP Site 1167 includes a thin mudstone horizon at 217 mbsf with calcareous nannoplankton not presently found in Prydz Bay (Shipboard Scientiﬁc Party, 2001b), suggesting warmer conditions at about 1.1 Ma (Pospical 2004; Lavelle, personal communication, 2001). These occurrences indicate warmer episodes when the Amery Ice Shelf edge retreated several hundred kilometres inland from its present position, and warmer water intruded Prydz Bay. 5.4.5. Prydz Bay Summary Seismic interpretation and drilling data reveal that the glaciation of Prydz Bay started in the latest Eocene. At that time, Prydz Bay was occupied by a ﬂuviodeltaic plain covered with stunted cool-temperate vegetation. Rivers ﬂowing through the Lambert Graben were fed by glaciers in the hinterland. The sea transgressed across the plain and ﬂoating ice delivered dropstones to the shallow embayment. The embayment became glaciated in the early Oligocene, with ice-sheet-scale glaciers depositing subglacial till, and glacimarine diamicts when the ice was not grounded. The ice was probably wet-based. In the early Miocene, a large temperate to polythermal ice sheet advanced and retreated across the embayment, supplying large quantities of detritus to the continental rise, where the detritus was deposited in large mounds. The mid-Miocene was marked by the start of a cooling trend and the development of a thicker, colder and more erosive ice sheet. Shelf overdeepening began, but progressively less detritus was delivered to the continental rise. In the early Pliocene, ice ﬂow became concentrated on the western side of the bay in an ice stream that deposited sediment in a trough mouth fan. During warm phases, open-water extended landward as far as the

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northern Prince Charles Mountains. Ice volumes and depths of erosion reached a peak in the mid Pleistocene and the cold, polar ice sheet was established. The Amery Ice Shelf no longer grounded at the shelf edge in Prydz Channel during glacial episodes.

5.5. Weddell Sea (Y. Kristoffersen and W. Jokat) The principal features in the Weddell Sea sector relevant to resolving the Antarctic paleoclimate and paleoceanographic history are prograding wedges of glacigenic sediments along the entire margin, a major troughmouth fan (Crary Fan), and numerous sediment drifts on the slope and in the deep basin, particularly along the western and northwestern side of the Antarctic Peninsula. Ice sheet ﬂow-line patterns suggest that the continental margin of the eastern and southern Weddell Sea east of 451W receives drainage from the EAIS, whereas the continental margin west of 451W receives drainage from the WAIS (Fig. WS-1).

Figure WS-1: Track lines of multichannel seismic data in the Weddell Sea and locations of ODP drill sites. Bathymetry after Schenke et al. (1998). Areal extent of proximal and distal deposits of the Crary Trough Mouth Fan are outlined by light brown shaded area. True extent of sediment drifts in the southwestern and western Weddell Sea is poorly deﬁned due to lack of data coverage. CTMF, Crary Trough Mouth Fan; EE, Explora Escarpment.

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5.5.1. The Regional Seismic and Geologic Database Modern geophysical data in the Weddell Sea comprise about 45,000 km of MCS lines from surveys principally by German, Norwegian and Russian research institutions since 1976 (Fig. WS-1). ODP drilled four sites in the Weddell Sea during ODP Leg 113, and ODP Site 693 on the Dronning Maud Land continental slope has been the most useful for stratigraphic calibration (Figs. WS-1 and WS-2). Prior to the deep drilling, stratigraphic studies in the region were conducted by Elverhøi and Maisey (1983), Hinz and Krause (1982), Hinz and Block (1984), Haugland et al. (1985) and Hinz and Kristoffersen (1987). Correlations with ODP Site 693 were made by Miller et al. (1990), Kuvaas and Kristoffersen (1991), Moons et al. (1992), Michels et al. (2002) and most extensively by Rogenhagen et al. (2004) (Fig. WS-2).

5.5.2. Acoustic Stratigraphy of the Shelf/Slope/Rise Environment-Spatial and Temporal Characteristics The continental shelf of the Weddell Sea is characterized by a prograding wedge of glacigenic sediments more than 1 km thick below the shelf edge (Fig. WS-3). The wedge downlaps onto older units, which are characterized by rather uniform thickness in the down-slope direction (Fig. WS-3a). Wedge deposition is a ﬁrst order result of massive transport of unsorted texturally immature sediments by advance of a grounded ice sheet to the shelf edge (Barker et al., 1998). The acoustic response of coarse sediment in proximal positions below the shelf and uppermost slope is one of discontinuous reﬂection events. Continuity and deﬁnition of acoustic stratiﬁcation improve in the down-slope direction as a result of progressive sorting and increased relative abundance of ﬁner material. The shelf edge may appear rectilinear, but the three-dimensional wedge architecture in the eastern Weddell Sea reveals an amalgation of adjacent small discrete cones of glacial sediments sourced by smaller ice streams (Kristoffersen et al., 2000). The spectrum of cones reﬂects broad scale expansion of the EAIS, but adjacent cones may or may not be coeval. Topsets of the prograding wedge are generally truncated at the seabed. Shelf aggradation is indicated in the southern Weddell Sea west of the Crary Trough, but the vast shelf area west of 451W has not been accessible for seismic surveys (Fig. WS-1). The maximum thickness of prograding units below the mouth of the Crary Trough and also below the shelf north of Lyddan Ice Rise is more than 3 km (Rogenhagen et al., 2004). ODP Site 693 (Fig. WS-2) provides local

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Figure WS-2: Stratigraphic summary column. Modiﬁed from Rogenhagen et al. (2004).

Figure WS-3: (a) Seismic line AWI-90110 across the Dronning Maud Land margin showing the prograding wedge (modiﬁed from Michels et al., 2002). (b) Seismic line AWI-97051 across the Larsen Shelf and Slope, showing the prograding shelf and sediment drift on the lower continental slope (modiﬁed from Michels et al., 2001). Proﬁle locations are in Fig. WS-1.

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Figure WS-4: Seismic line NARE-8517 across the prograding Crary Trough Mouth Fan (modiﬁed from Kuvaas and Kristoffersen, 1991). Proﬁle location is in Fig. WS-1. calibration of the acoustic section (Miller et al., 1990), but regional extrapolations are inhibited along-slope by numerous canyons, and are inhibited down-slope by the steep Explora Escarpment (Fig. WS-1). The Crary Fan, a regional feature at the mouth of the Crary Trough, is associated with large channel/levee complexes, which extend up to 1,000 km to the north into the basin (Figs. WS-1, WS-4 and Foldout WS-1). Initial fan evolution is correlated with the resumption of sediment deposition above an Albian-early Oligocene hiatus at ODP Site 693 (Reﬂector W4). Sediment drifts are common within the Neogene stratigraphic interval along the continental slope (Fig. WS-1) in the western Weddell Sea (Michels et al., 2001; Maldonado et al., 2005).

5.5.3. The Weddell Sea Pre-Ice-Sheet Depositional Environment Acoustic stratigraphic information on the shelf is limited to subbottom depths comparable to the local water depth due to severe multiple reﬂections (e.g. Fig. WS-3). The pre-ice-sheet Cenozoic shelf edge was more than 10 km landward of its current position along the Dronning Maud Land continental margin (Kristoffersen et al., 2000), and 70 km to the south (Fig. WS-4) in the southern Weddell Sea (Kuvaas and Kristoffersen, 1991). The shoreward

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shift in the western Weddell Sea is unknown. The deeper strata below the continental slope (i.e. below W4) appear unstructured throughout. The older sediments are thickest (5–8 km) below the Larsen Shelf in the western Weddell Sea (Rogenhagen and Jokat, 2000), and may be up to 15 km thick along the front of the Ronne and Filchner Ice shelves (Leitchenkov and Kudryavtzev, 2000). In the central Weddell Sea Basin, the pre-Oligocene section of inferred turbidites is more than 1 km thick, and thins by basal onlap towards the margins to less than 0.5 km (Rogenhagen et al., 2004). High seasonal variations in sea-surface temperatures and a well-developed seasonal thermocline characterized the early Paleogene Weddell Sea (Kennett and Barker, 1990). On Maud Rise, siliceous biogenic facies began to replace carbonate facies during the latest Eocene-earliest Oligocene (Kennett and Barker, 1990). A possible early Cenozoic seaway between East and West Antarctica could have been up to 700 m deep, and may have persisted into the Oligocene if no WAIS was present (Lawver and Gahagan, 2003). At ODP Site 693 on the middle continental slope, middle lower Oligocene and younger glacial sediments are separated by a hiatus from Albian radiolarian diatomite and claystones (i.e. Reﬂector W4). The unconformity may represent non-deposition and/or mild erosion (Kennett and Barker, 1990). 5.5.4. Change from Non-Glacial to Glacial Conditions Sediment ﬂuxes on high latitude continental margins are closely connected to climate extremes. In the Weddell Sea, environmental change is manifested by a basin-wide change in acoustic character within the sedimentary section (Reﬂector W4) at about 1 s TWT below the sea bed (Rogenhagen et al., 2004). Younger deposits in the basin have ﬁnely laminated continuous acoustic stratiﬁcation, and geometries on the slope are in the form of channel/levee complexes over a wide range of spatial scales. The change in depositional environment is interpreted to have originated from an increased sediment ﬂux, caused by increased erosion of the continent and increased down-slope transport. At ODP Site 693 on the middle continental slope, the acoustic change correlates stratigraphically with resumed preservation of lower Oligocene sediments. The deposits include rounded dropstones in lower Oligocene (32–33 Ma) diatom muds, a signal of the ﬁrst presence of glaciers on the adjacent parts of East Antarctica (Kennett and Barker, 1990). Subsequent early Miocene sedimentation rates at this site were low (7 m/Ma). A more dramatic change in sediment ﬂux to the margin is documented by a threefold increase in sedimentation rate (to 24 m/m.y.), when sedimentation

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resumed following a hiatus that spanned the middle Miocene. Increased sediment input is related to expansion of ice on the East Antarctic continent. The hiatus at ODP Site 693 correlates with a regional acoustic reﬂection event (W5) identiﬁed below the continental slope and rise along the entire Weddell Sea margin (Rogenhagen et al., 2004). Shelf progradation accelerated dramatically along the eastern and western margins of the Weddell Sea (Fig. WS-3), with grounded ice extending to the shelf edge in the late Miocene (Michels et al., 2001; Michels et al., 2002). A range of contourite drifts formed on the slope and rise in the northwestern Weddell Sea (Michels et al., 2001; Maldonado et al., 2005). Sedimentation rates at ODP Site 693 reached 60 m/m.y. in the early Pliocene, and subsequent Quaternary sedimentation rates were reduced to 16 m/m.y. (Gersonde et al., 1990). Sediment input to the margin in the southeastern Weddell Sea was focused toward a trough mouth fan. The Crary Fan began to expand at the time of change to a glacial environment (above Reﬂector W4, Fig. WS-4 and Foldout WS-1), and major channel/levee complexes evolved in three phases. The last of these three phases (Reﬂector W5, Fig. WS-4 and Foldout WS-1) was from the late Miocene on (Kuvaas and Kristoffersen, 1991; Moons et al., 1992). 5.5.5. The Glacial/Interglacial Environment The change from a glacial to an interglacial environment was associated with major changes in sediment ﬂux. Average sediment deposition on the eastern Weddell Sea margin (101W) during the last two climatic cycles (300 k.y.) varies from 5 cm/k.y. on the upper slope to over 1 cm/k.y. on the lower slope (Grobe and Mackensen, 1992). Sedimentation was most rapid during the beginning of interglacials, with rates on the middle slope four to ﬁve times higher than during glacials. We note, however, that the grounded EAIS only reached the mid-shelf in this area during the LGM (Kristoffersen et al., 2000). Sediment input in the southern Weddell Sea was focused in the Crary Trough Mouth Fan (Figs. WS-1, WS-4 and Foldout WS-1). The fan comprises large channel-levees on the ﬂanks of deep-water channels, such as the Cold Water Channel and the Deutschland Channel (Foldout WS-1). Grounded ice reached the shelf edge at the trough mouth during the last glaciation (Bentley and Anderson, 1998), and deposition on the levees (in water depths of 2,000–3,000 m) ranged from 100–200 cm/k.y. during the LGM to a few cm/k.y. during the present interglacial (Weber et al., 1994). Episodic sediment transport into the basin also occurred by mass ﬂows during interglacials, probably as partial collapse of the deposits on the upper

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continental slope. A 90-m-thick sandy turbidite unit was deposited within 0.5 m.y. during the early Gilbert Chron (4.8 Ma) at ODP Site 694 (Fig. WS-1), and may be the distal expression of mass wasting events on the continental slope in the southwestern Weddell Sea (Shipboard Scientiﬁc Party, 1988). Also, major early Pliocene drawdown of East Antarctic ice is postulated to have triggered extensive mass ﬂows that originated from the Crary Trough Mouth Fan (Bart et al., 1999). In the western Weddell Sea, upper Miocene and younger sediments (above Reﬂector W5) are mostly drift deposits that reach a thickness of more than 1 km below the middle slope, seaward of the Larsen Shelf (Rogenhagen and Jokat, 2000; Michels et al., 2001; Maldonado et al., 2005). Present and past bottom currents circulated in nearly the opposite direction to channel transport, and cross-channel ﬂow was in the same direction as the Coriolis force acting on down-slope turbidity currents in the southern Weddell Sea. Sediments scavenged from turbid channel ﬂow by cross-channel bottom currents sourced the benthic boundary layer and enhanced formation of sediment drifts along the western and northern Weddell Basin. The actual drift distribution was mainly controlled by the physiography of the basin and bottom current ﬂow directions (Maldonado et al., 2005). These drifts represent a storehouse of paleoceanographic and climatic proxies not yet sampled by scientiﬁc drilling. 5.5.6. Continental Margin Sediments and Ice-Sheet History The mass balance of the EAIS, the nature of the substratum and the continental topography, particularly in the coastal region, determine sediment input to the continental margin. Enhanced input of sediments to the continental margin at ODP Site 693 in the eastern Weddell Sea and development of a prograding wedge started in the latest Miocene and peaked during the earliest Pliocene (Gersonde et al., 1990). The seismic tie between ODP Site 693 and the southern Weddell Sea is uncertain, but Kuvaas and Kristoffersen (1991) suggest that fan development started in the southern Weddell Sea by the early Oligocene (above Reﬂector W4, Fig. WS-4 and Foldout WS-1), and that about two-thirds of the sediment thickness at the mouth of the present Crary Trough was already in place by the late Miocene (i.e. below Reﬂector W5). Channel-levee complexes have migrated eastward on the Crary Trough Mouth Fan, and late Miocene and younger deposition constructed a third major channel-levee complex and deposited about 1 km of sediments below the trough mouth (Fig. WS-4 and Foldout WS-1). These age relations imply that the principal input of sediments from East

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Antarctica to the Weddell Sea margin from the early Oligocene to the late Miocene originated from a glaciated interior of the continent via the Crary Trough, and that there was effectively no input along the Dronning Maud Land margin. At this point, the signiﬁcance of a local thickness maximum of glacial sediments north of Lyddan Ice Rise (Rogenhagen et al., 2004) is unclear. The EAIS expanded to the Dronning Maud Land margin during the latest Miocene–earliest Pliocene and formed a prograding wedge below the continental shelf and slope. Sea ice cover has prevented acquisition of the seismic data from west of 451W and north of the Ronne Ice Shelf (Fig. WS-1) needed to study the depositional geometries of sediments originating from the catchment area of the WAIS. Data from this area also are needed to study the relation between eastern and western sediment source regions. Moraine complexes on the shelf in the eastern Weddell Sea suggest that the EAIS was grounded on the mid-shelf and did not reach the shelf edge during the LGM (Kristoffersen et al., 2000), except at the mouth of the Crary Trough (Bentley and Anderson, 1998).

5.6. Antarctic Peninsula (R. Larter) Cenozoic tectonic processes have diversely affected the Antarctic Peninsula region and its climate record. Hence, we separately discuss four main subregions.

5.6.1. The Eastern Margin This subregion includes the Weddell Sea margin of the Antarctic Peninsula and Larsen Basin (Fig. AP-1). Persistent sea ice covers the region (Gloersen et al., 1992), hence relatively few research cruises have been conducted here. Macdonald et al. (1988) used regional geologic and aerogeophysical data to infer that a large Mesozoic–Cenozoic sedimentary basin extends B700 km south from James Ross Island. Four main seismic stratigraphic units are identiﬁed from SCS reﬂection data on the shelf and upper slope (Anderson et al., 1992; Sloan et al., 1995; Fig. AP-2): Unit 4: acoustic basement interpreted as Jurassic and younger volcanic rocks; Unit 3: seaward-dipping reﬂections interpreted as Late Cretaceous to Oligocene marine shelf deposits, the older part of which are coeval with those on nearby Seymour Island; Unit 2: prograding sequences with truncated foresets that downlap onto Unit 3, and that are thought to have been deposited by multiple advances of grounded ice across the shelf in the Miocene and early Pliocene;

Figure AP-1: (a) Track lines of multichannel (thick lines) and single-channel (thin lines) seismic data in the Antarctic Peninsula region, DSDP and ODP drill sites (ﬁlled circles, annotated with site numbers) and SHALDRIL sites (open squares). SHALDRIL sites are only marked where either W10 m subseaﬂoor penetration was achieved or pre-quaternary sediments were recovered. Bathymetric contours are at 1,000 m intervals down to 4,000 m, with 500 m contours included locally on the shelf. Bathymetric data are based on Smith and Sandwell (1997) east of the Peninsula and north of 621S. Bathymetry for most of the area west of the Peninsula is from Rebesco et al. (1998), the exception being the 500 m contours in the southern Bellingshausen Sea, which are from O´ Cofaigh et al. (2005). To the east of the Peninsula, 500 m contours south of James Ross Island (JRI) are based on ﬁgures in Evans et al. (2005), and the 500 m contour north of James Ross Island is based on multibeam echo sounding data maps produced from data collected on RV Nathaniel B. Palmer Cruises 0003 and 0107. Bold line is location of seismic proﬁle in Fig. AP-2. SOM: South Orkney Microcontinent; JB: Jane Basin; PB: Powell Basin; BS: Bransﬁeld Strait; RT: Robertson Trough. (b) Expanded map of the northern Antarctic Peninsula with SHALDRIL site numbers labelled (I/II indicates ﬁrst/second cruise). Same bathymetry contours shown as in (a), except 500 m contours around JRI omitted. SSI: South Shetland Islands.

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Unit 1: aggrading reﬂections interpreted as deposits from ﬂuctuating dynamic ice sheets in the Pliocene and Pleistocene. MCS reﬂection data show at least 8 km of sediment at the base of the continental slope, overlying likely Early Cretaceous age basement (Barker and Lonsdale, 1991). The northern continental slope has plastered contourite drift deposits up to 900 m thick, thought to have been deposited by northﬂowing glacially inﬂuenced bottom currents. Pudsey (2002) suggests that drift deposition began in the early Miocene at the onset of bottom water ﬂow, or in the latest Miocene at the onset of volumnious glacially derived sediment supply to the western Weddell Sea. Late Quaternary shelf sediments have been sampled by seaﬂoor coring (e.g. Domack et al., 2001a, b, 2005; Pudsey and Evans, 2001; Pudsey et al., 2001; Brachfeld et al., 2003; Evans et al., 2005). These researchers infer that grounded ice converged into major ice streams and advanced to the shelf edge during the LGM, that the Prince Gustav Sound Ice Shelf collapsed and reformed in the mid-Holocene, and that the recent collapse of the Larsen B Ice Shelf is unprecedented during the Holocene. Recent drilling by the SHALDRIL project has obtained the ﬁrst samples from older sequences (Shipboard Scientiﬁc Party, 2005, 2006; Anderson et al., 2006, 2007).

Figure AP-2: Single-channel seismic line across the eastern margin of the Antarctic Peninsula at 65115uS, collected on RV Nathaniel B. Palmer in 1993. Line drawing interpretation shows the internal geometries and boundaries (dashed lines) between the main seismic units. Vertical exaggeration at the seaﬂoor is 82:1. Adapted from Sloan et al. (1995). Line location is shown in Fig. AP-1.

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Sequences drilled by SHALDRIL are principally shallow marine sands, sandy and silty muds, and pebbly muddy sands from neritic environments, with mollusc shells distributed throughout the cores. The cores are dated as: late Eocene to early Oligocene, Oligocene, middle Miocene, early Pliocene and Holocene. 5.6.2. The South Orkney Islands Region This subregion includes the South Orkney Microcontinent (SOM) and the adjacent deep-water Jane and Powell basins (Fig. AP-1). The SOM extends about 350 km from east to west and 250 km from north to south, and is underlain by Mesozoic metamorphic and sedimentary rocks (Thomson, 1981; Dalziel, 1984). Offshore, the SOM includes four Cenozoic sedimentary basins (King and Barker, 1988) with up to 5 km of sediment (Harrington et al., 1972; Busetti et al., 2001, 2002). Powell Basin (up to 3,600 m deep) formed as the SOM rifted and drifted away from the tip of the Antarctic Peninsula in late Eocene to late Oligocene time (King and Barker, 1988; Lawver et al., 1994; Coren et al., 1997; Eagles and Livermore, 2002). Opening of Jane Basin (up to 3,300 m deep) probably began slightly later (Lawver et al., 1991, 1994), and may have continued until the middle Miocene according to Maldonado et al. (1998). From SCS data on the SOM, King and Barker (1988) deﬁned pre-rift, synrift, and post-rift units. The post-rift sediments are less than 1 km thick (e.g. Busetti et al., 2001, 2002) and were drilled at ODP Site 695 (1,300 m water depth) and ODP Site 696 (600 m water depth) (Fig. AP-3). They comprise Oligocene or early Miocene to Quaternary terrigenous sediments, with rare coarse-grained IRD until the late Miocene (B8.7 Ma) and common IRD thereafter. Middle Miocene to Quaternary sediments are hemipelagic and diatomaceous muds and oozes (Barker et al., 1988a, b). ODP Site 696 also sampled syn-rift Eocene sandy mudstones (Sequence 2) that have nannofossil assemblages and clay minerals suggesting a relatively warm climate, and palynoﬂora indicating temperate beech forests and ferns on West Antarctica. Drilling results suggest intermittent glaciation with little sea ice during most of the Miocene and a persistent ice cap to sea-level on West Antarctica since the late Miocene. Herron and Anderson (1990) place the maximum late Quaternary grounding line advance at the 300 m isobath, and consider openmarine conditions to have existed over the SOM since 6,000 y. B.P. based on SCS and seaﬂoor core data. In Powell Basin, post-early-rift sediments are up to 3 km thick. King et al. (1997) identiﬁed two seismic units with low reﬂectivity below and high

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Figure AP-3: Part of seismic line AMG845-18, showing the setting of ODP Site 696 in relation to the seismic units (S1–S3) described by King and Barker (1988). Vertical exaggeration at the seaﬂoor is 3.3:1. Adapted from Barker et al., 1988a. ODP Site 696 location shown in Fig. AP-1. reﬂectivity above. They interpret the change as recording the onset of glacial–interglacial cyclicity in the supply of coarse detritus to the basin in the late Miocene. A similar upward change in reﬂectivity is observed in Jane Basin (Maldonado et al., 1998). The reﬂectivity change may also be due to silica diagenesis (e.g. Lonsdale, 1990; Volpi et al., 2003). Maldonado et al. (2006) identify ﬁve seismic units in Jane and neighbouring ocean basins, and relate changes in seismic characteristics to variations in bottom water ﬂow since the middle Miocene. ODP Site 697 was drilled in Jane Basin (Fig. AP-1) to B323 mbsf, and recovered mainly early Pliocene and younger hemipelagic sediments with IRD throughout; however, IRD is abundant only near the base of the sequence (Barker et al., 1988a, b). Other seismic studies of Powell Basin (e.g. Kavoun and Vinnikovskaya, 1994; Coren et al., 1997; Viseras and Maldonado, 1999) focus on the post-early Oligocene rift history of the basin, but are limited in paleoclimate interpretations by the lack of drilling data.

5.6.3. The South Shetland Islands Region This subregion includes the Bransﬁeld Strait and the continental margin around the South Shetland Islands (Fig. AP-1). Bransﬁeld Strait is a 2,000 m deep rift basin that is actively extending at 7 mm/y. (Dietrich et al., 2004). The time of initial extension and the oldest age of basin sediments are uncertain, but may be 4 Ma (Barker and Dalziel, 1983) or B6 Ma (Larter and Barker, 1991a). Gamboˆa and Maldonado (1990) speculate that

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Bransﬁeld Strait may have opened earlier, during the early Miocene, and have been continuous with mid-shelf basins to the southwest. From MCS data, Gamboˆa and Maldonado (1990) identify ‘rift’ and ‘drift’ sequences in Bransﬁeld Strait. ‘Drift’ sequences prograde the shelf and are about 1 km thick. In SCS data, Jeffers and Anderson (1990) deﬁne four glacio-eustatic sequences within the ‘drift’ sequences, and interpret all four sequences as being younger than 3 Ma. Prieto et al. (1999) use different SCS data to deﬁne eight seismic units that comprise interﬁngering slope and basinal deposits within the ‘drift’ sequences. They interpret the slope units as having been deposited directly from grounded ice during glacial periods, and interpret the basinal units as having been deposited by mass ﬂow processes during deglaciations and interglacial periods. From SCS data, Banﬁeld and Anderson (1995) identify sediment mound features that they infer to be glacial grounding moraines in up to 1,000 m water depth on the southeastern ﬂank of Bransﬁeld Strait. They speculate that mounds at B700 m depth mark the maximum advance during the LGM. Deep troughs with mega-scale lineations are incised into the shelf and are interpreted as the paths of palaeo-ice streams (Banﬁeld and Anderson, 1995; Canals et al., 2002). MCS proﬁles across the continental slope NW of the South Shetland Islands reveal a forearc basin, with more than 1.5 km sediments, that is bounded to the NW by a small accretionary prism (Maldonado et al., 1994a, b). The prism overthrusts trench-ﬁll sediments that may have been deposited rapidly and are up to 1 km thick (Maldonado et al., 1994a, b; Kim et al., 1995). Other seismic-reﬂection surveys have been done in the region by British, Polish, German, Spanish, US, Italian, Chinese and Korean research groups, but published results focus on the tectonic evolution of the region and tectonic processes (Barker, 1976; Guterch et al., 1985; GRAPE Team, 1990; Acosta et al., 1992; Henriet et al., 1992; Grad et al., 1993; Barker and Austin, 1994, 1998; Bochu et al., 1995; Gra`cia et al., 1996; Jin et al., 1996; Jin and Kim, 1998; Prieto et al., 1998; Jin et al., 2002), on gas hydrates (Lodolo et al., 1993, 2002; Tinivella et al., 1998, 2002; Jin et al., 2003), and on a large submarine slide (Imbo et al., 2003). Swath bathymetry data exist over most deep-water parts (W1,000 m) of Bransﬁeld Strait (Lawver et al., 1996; Gra`cia et al., 1997). Along Boyd Strait, ‘bundle structures’ and mega-scale glacial lineations occur and conﬁrm palaeo-ice stream ﬂow during glacial periods (Canals et al., 2000; COHIMAR/SEDANO Scientiﬁc Party, 2003). MCS data along outer Boyd Strait reveal glacial progradation of the margin (Maldonado et al., 1994a, b), and deep-tow boomer proﬁles reveal a glacier grounding zone wedge near the shelf edge (Vanneste and Larter, 1995).

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The only scientiﬁc drilling in the South Shetland Islands region was done at SHALDRIL-I Site 1 in Maxwell Bay (Fig. AP-1), where an expanded sequence of Holocene diatomaceous muds, B105 m thick, overlying a clayrich diamicton was sampled (Shipboard Scientiﬁc Party, 2005). 5.6.4. The Paciﬁc Margin The Paciﬁc margin includes the region southwest from the South Shetland Islands to 701S and 801W. This is a former active margin where ridge-crest segments were progressively subducted (Larter and Barker, 1991a; Henriet et al., 1992), followed by 1–4 m.y. of uplift and then long-term subsidence (Larter and Barker, 1989, 1991b; Anderson et al., 1990; Gamboˆa and Maldonado, 1990; Bart and Anderson, 1995, 1996, 2000; Larter et al., 1997). Seismic proﬁles show that outer shelf sequences are separated from NW-SE trending mid-shelf basins by the Mid-shelf High (MSH) (Kimura, 1982; Anderson et al., 1990; Gamboˆa and Maldonado, 1990; Larter et al., 1997). Several research groups have conducted seismic studies in the area (Kimura, 1982; Larter and Barker, 1989, 1991b; Anderson et al., 1990; Gamboˆa and Maldonado, 1990; Henriet et al., 1992; Bart and Anderson, 1995, 1996, 2000; McGinnis and Hayes, 1995; Rebesco et al., 1996, 1997, 2002, 2006; McGinnis et al., 1997; Larter et al., 1997; Jin et al., 2002; Jabaloy et al., 2003; Herna´ndez-Molina et al., 2006a). Cenozoic sequences have been drilled at DSDP Site 325 (Hollister et al., 1976) and at multiple sites during ODP Leg 178 (Barker et al., 1999). Evidence of Oligocene glaciation exists on King George Island (Birkenmajer, 1991; Dingle and Lavelle, 1998; Troedson and Smellie, 2002), but offshore the ﬁrst ice sheets on the Paciﬁc margin are inferred from early Miocene IRD at DSDP Site 325 (Fig. AP-1). The oldest sediments from ODP Leg 178 are drift deposits at ODP Site 1095, dated at 9.6 Ma, where all cores show glacial inﬂuence and sedimentation rates decrease steadily from the late Miocene to the Quaternary. From seismic studies, Rebesco et al. (1996, 1997, 2002) suggest that sediment drift deposition began around the middle Miocene, with most growth in the late Miocene. However, Herna´ndez-Molina et al. (2004, 2006b) describe a buried sediment drift of early Miocene age (Fig. AP-1). Uenzelmann-Neben (2006) also interprets depositional patterns of early Miocene continental rise sediments as reﬂecting bottom current inﬂuence, but infers a different ﬂow direction. A regular supply of both glacially derived terrigenous sediments and

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interglacial biogenic sediments has reached the continental rise since at least the middle Miocene. The outer continental shelf is underlain by depositional sequences that thicken seaward (Larter and Barker, 1989, 1991b; Anderson et al., 1990; Gamboˆa and Maldonado, 1990; Larter and Cunningham, 1993; Bart and Anderson, 1995; Table AP-1). Aggrading sequences (S3) without a distinct paleo-shelf edge (PSE) are unconformably overlain by prograding sequences (S2, S1) with an abrupt PSE (Foldout AP-1). The change occurs at the S3/S2 boundary of Larter and Barker (1989, 1991b) and is observed all along the Paciﬁc margin (Anderson et al., 1990; Bart and Anderson, 1995, Larter et al., 1997; Jin et al., 2002) and west along the margin of the Bellingshausen and Amundsen Seas (Nitsche et al., 1997). Drilling at ODP Sites 1097 and 1103, although with very poor recovery, suggests that the change occurred between 8 and 6 Ma (Iwai and Winter, 2002; Bart et al., 2005). Cores from S3 are diamictons with interbedded mudstones and graded sandstones, interpreted by Eyles et al. (2001) as continental slope deposits, although seismic proﬁles suggest a palaeo-shelf to slope transition further offshore. Cores from S1 and the upper part of S2 show abundant evidence for having been deposited subglacially (Eyles et al., 2001). Larter and Barker (1989, 1991b) and Larter et al. (1997) interpreted the S3/S2 boundary as representing the onset of frequent advances of grounded ice to the palaeo-shelf edge. However, Bart and Anderson (1995) and Bart et al. (2005, 2007) suggest that palaeo-ice streams cut erosional troughs within S3 and hence existed earlier than the S3/S2 boundary. Although equivocal, the boundary could represent a change in the typical extent of glacial advances, in the dynamic behaviour of ice sheets that advanced onto the shelf, or in the way the ice transported sediments (Larter et al., 1997; Herna´ndez-Molina et al., 2006a). Going up-section above the S3/S2 boundary, foreset stratal dips generally increase and PSE progradation in individual sequences decreases (Foldout AP-1). The regional S2/S1 boundary within the upper sedimentary section may have been produced by ice-sheet erosion during glacial periods and lower sea levels after the Late Pliocene increase in the volume of Northern Hemisphere ice sheets (Larter and Barker, 1989, 1991b). By seismic correlation across W100 km to ODP Site 1101, Rebesco et al. (2006) estimate an age of B3 Ma for a boundary that they identify as S2/S1. However, this boundary marks a change in stratal geometry that is generally characteristic of S3/S2, and Larter (2007) suggests that it probably corresponds to this earlier unconformity.

Table AP-1: Summary of stratigraphic schemes used in previous publications to describe the late Miocene to recent depositional sequences on the Paciﬁc margin of the Antarctic Peninsula.

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Mid-shelf basins contain up to 2 km of sediment in a broad synform that is truncated at shallow depth beneath the seaﬂoor (Kimura, 1982; Anderson et al., 1990; Gamboˆa and Maldonado, 1990; Larter et al., 1997), making the succession accessible to shallow drilling. However, these sediments have not been sampled, except for the thin Quaternary cover. Basin sediments are probably all Tertiary and may be as young as early Miocene off Adelaide Island, and middle Miocene off Anvers Island (Larter et al., 1997). The inner shelf is mostly shallower than 200 m, but has deep troughs, such as the Palmer Deep where ODP Sites 1098 and 1099 were drilled in 1,400 m of water. Holocene successions B47 and B108 m thick were recovered at ODP Sites 1098 and 1099, respectively (Shipboard Scientiﬁc Party, 1999; Domack et al., 2001; Ishman and Sperling, 2002; Leventer et al., 2002; Shevenell and Kennett, 2002). Swath bathymetric data show seaﬂoor features of subglacial origin on the shelf (O´ Cofaigh et al., 2002; Dowdeswell et al., 2004; Amblas et al., 2006), and conﬁrm that a grounded ice sheet with ice streams extended to the shelf edge during the LGM (Pudsey et al., 1994; Larter and Vanneste, 1995). Seaﬂoor core data indicate that retreat of the ice sheet from the outer and middle shelf after the LGM occurred between 18,500 and 13,000 cal. y. B.P. (Pudsey et al., 1994; Heroy and Anderson, 2005).

5.7. Other Sectors of the Antarctic Continental Margin Other sectors of the Antarctic margin, than the ﬁve ANTOSTRAT project working areas discussed above, have been studied more fully than previously during the past 5–7 years, in a time of renewed interest in MCS studies of the continental margin. These regions include similar geomorphic and stratigraphic features. Although we do not include them in the discussion below due to space considerations, we include here representative citations to some of the studies published for the Bellingshausen and Amundsen seas (Hollister et al., 1976; Tucholke and Houtz, 1976; Tucholke, 1977; Kimura, 1982; Yamaguchi et al., 1988; Cunningham et al., 1994, 2002; Gohl et al., 1997, 2007; Nitsche et al., 1997, 2000; Wellner et al., 2001, 2006; Lowe and Anderson, 2002; O´ Cofaigh et al., 2005; Dowdeswell et al., 2006; Evans et al., 2006; Scheuer et al., 2006a, b; Larter et al., 2007; Uenzelmann-Neben et al., 2007) and for offshore areas of Queen Maud Land and Enderby Land (Kuvaas et al., 2004a, b, 2005; Hinz et al., 2005; Stagg et al., 2004a, b; Leitchenkov et al., 2007; Solli et al., 2007a, b, c).

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5.8. Discussion Our discussion focuses on integrating key observations and inferences from the Antarctic continental margin to summarize a Cenozoic glacial history from the stratigraphic record. We recognize the limitations of the seismic and drilling data sets used. Seismic data image regional and up-section changes, but only provide a relative history of inferred events and processes. Drilling data provide a deﬁned stratigraphic history, but one that is valid only at the local drilling site. There is an extensive published literature that illustrates the large spatial and temporal variability of features and processes at various scales, depending on the resolution and extent of data analysed. We selectively discuss the widespread seismic stratigraphic variations and the long- and short-period geologic transitions in cores, all of which point to a varied history of non-glacial and glacial events on the continental margin. The discussion is necessarily condensed due to the great breadth of the topics and the limited space herein. We abbreviate offshore geographic regions as: Ross Sea (RS), Antarctic Peninsula (AP), Weddell Sea (WS), Prydz Bay (PB) and Wilkes Land (WL), and preﬁx these abbreviations with east (E) and west (W), such as western Ross Sea (WRS) (Fig. I-1).

5.8.1. Regional Seismic Stratigraphic Variations: Similarities and Differences Stratigraphers have long recognized sedimentary sequences and bounding regional unconformities in seismic-reﬂection data across the Antarctic continental margin (e.g. Hinz and Block, 1984; Wannesson et al., 1985; Hinz and Kristoffersen, 1987; Cooper et al., 1995). These sequences, principally of inferred Cenozoic age, commonly have similar seismic geometries around Antarctica (e.g. Cooper et al., 1991b; Anderson, 1999). Some geometries are unique to polar continental margins, whereas other geometries are like those of low-latitude non-polar margins (e.g. Hinz and Block, 1984; Bartek and Anderson, 1991; Table D-1). A uniﬁed circum-Antarctic seismic stratigraphy does not exist, but an International effort to compile one via the CASP project is in progress (Davey and Cooper, 2007). Numerous separate and sometimes different seismic stratigraphies exist for various localities. Many seismic features have been cited to suggest intermittent ice on the Antarctic margin (Table D-1), but those giving the strongest evidence for ice are the regional seismic unconformities, the broad erosional troughs and depositional banks on the continental shelf, and the large-scale fans and

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Table D1-1: Common large-scale geomorphic and seismic stratigraphic features of the Antarctic continental margin (listed by location and decreasing inferred age). Feature Continental shelf A. Deep regional seismic unconformities on the shelf (lower B1/3 of sedimentary section) B. Shallow regional seismic unconformities on the shelf (upper B2/3 of section) C. Prograded and aggraded sequences under the outer shelf D. Mound features with chaotic seismic facies E. Broad cross shelf troughs and adjacent banks F. Overdeepened and foredeepened continental shelf Continental slope G. Regional seismic unconformities H. Massive sediment fan on the slope at the outlet of a broad cross-shelf trough I. Steep slopes and migrating high-relief channels J. Variable size sediment fans on the upper slope at mouths of seaﬂoor and buried troughs. Continental rise K. Regional seismic unconformities

Regions

Processa

Timingb

PB, RS

Eustacy

Cretaceous to late Oligocene (?)

All

Ice-sheet erosion

late Oligocene (?) and younger

All

Oligocene and younger

All

Sediment carried to the shelf by glaciers Ice-sheet deposition

All

Ice-stream erosion

All

Ice-sheet erosion

mid-Miocene and younger

All

Bottom currents

PB, WS

Sediment deposited at shelf edge by one broad ice stream Bottom currents with coarse sediment Sediment deposited at shelf edge by multiple ice streams

Cretaceous and younger early Oligocene and younger (WS); early Pliocene and younger (PB) Oligocene and younger

All

WL, RS, AP

All

Bottom currents and diagenesis

late Oligocene and younger early Miocene and younger

late Miocene and younger

Cretaceous and younger

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Table D1-1: (Continued ). Feature

Regions

L. Large sediment drift features

All

M. An up-section landward shift of depocenters from rise to slope.

All

Processa Bottom currents and downslope sediment supply Reduction of sediment supply to rise

Timingb early Miocene and younger mid-Miocene and younger

a

The principal process is listed – others such as lithospheric loading, sediment loading, paleoceanographic processes, diagenesis, etc. are also commonly involved. b Initiation time of features varies in different regions.

prograding deposits on the continental slope at the mouths of the shelf troughs. These features are observed on all sediment-covered segments of the Antarctic margin, in East and West Antarctica, and principally occur in the upper part of the stratigraphic section. These features are increasingly common up-section, indicating more abundant glacial events more recently. The ubiquitous overdeepened and foredeepened depth proﬁle of the continental shelf is the ultimate evidence of sustained strong glacial erosion of the entire Antarctic margin. Ten Brink et al. (1995) modelled the geometries of the seaﬂoor and stratigraphic sections on the continental shelf and upper slope, incorporating lithospheric, glacial and eustatic processes in the models. They showed that multiple advances and retreats of grounded ice sheets across the continental shelf, coupled with redistribution of sediment from onshore and shelf areas to the continental slope, are required to match the observed geometries. Eustatic and paleoceanographic processes are important for sediment redistribution, especially on the continental slope and rise, but are not sufﬁcient by themselves to explain the shelf erosion and prograding geometries beneath the outer continental shelf. Bartek et al. (1991) illustrated that the stratal signatures of the Oligocene and younger sections in the eastern Ross Sea (i.e. unconformities and prograding sections) are similar to those of low-latitude non-polar margins, indicating that the Neogene stratal signature results from glacio-eustatic ﬂuctuations. Seismic data provide a relative history of increasing circum-Antarctic glacial events, as noted above, but drilling and seaﬂoor coring provide the only absolute age control and ground truth of glacial lithologies and processes. Deﬁnitive ages are limited due to the small number of cored sites, which in turn are biased toward sampling of shallow younger sections.

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The lithostratigraphic record from proximal drilling and dredging on the Antarctic margin ‘establishes’ a general history (at the drill sites) of no regional glaciers in Cretaceous and earlier times on the shelf (PB, RS, WL) and on the slope (WS). Evidence of glacial episodes is ﬁrst seen in the late Eocene to early Oligocene as diamicts from grounded ice on the shelf (PB, RS) and from glacial erratics on the slope (WS). Upper Oligocene glacial marine deposits are sampled on the shelf (RS) and slope (WS). Lower Miocene sections show increasing evidence of ice and deep-ocean currents, with IRD and sediment drifts on the rise (PB, WL, AP, WS?) and glacial marine sediments and diamicts on the shelf (RS). The middle Miocene has an increased glacial hemipelagic signature on the rise (PB), and glacial marine deposits on the shelf (RS). The late Miocene and early Pliocene were times of enhanced glacial activity, as recorded by: (1) shelf deposits of glacial diamicts (PB, AP) and glacial marine sediments (RS, AP); (2) glacial marine sediments on the slope (WS, WL); and (3) rise deposits of glacial hemipelagics (PB, WL, RS, AP) and turbidites (WL, WS, AP). Upper Neogene deposits are principally glacial on the shelf (PB, RS, AP), slope (PB, WS) and rise (PB, WL, RS, AP, WS). Based on recovered drill cores and Antarctic Peninsula coastal geology, the general history of events for East and West Antarctica is essentially the same since the early Oligocene. However, the middle Miocene and Oligocene history for offshore West Antarctica is based on only two Ross Sea DSDP cores (Hayes and Frakes, 1975) and two tentatively dated SHALDRIL cores from the western Weddell Sea (Anderson et al., 2006; Shipboard Scientiﬁc Party, 2006). The above general history at the sparse drill sites has been greatly expanded by many investigators who have traced seismic unconformities and seismic stratigraphic units from core sites and rare onshore sedimentary sections up to hundreds of km to infer ages and lithofacies for key stratigraphic features listed in Table D-1. The expanded seismic stratigraphic history includes inferences of the following (although uncertainties about these inferences are inherently greater than the uncertainties about ages at drillsites): A pre-ice-sheet to early-glacial period (Cretaceous to early Oligocene): On the inner and mid shelf, variable seismic facies (WRS, PB), narrow channel geometries (PB) and sea-ﬂoor dredged/cored rocks indicate the presence of pre-ice-sheet subaerial, ﬂuvial and shallow marine environments in the Cretaceous. These are unconformably overlain (WRS, PB) by upper Eocene to lower Oligocene early glacial sediments. This transition, from pre-ice-sheet to early glacial conditions, has not been sampled elsewhere on the shelf. Beneath the slope and upper rise, thick sediment sections are observed (all areas), but are not yet well imaged and mapped in most areas.

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Aggrading shelf period (Oligocene): On the shelf, stratal geometries in drilled Oligocene glaciomarine sections mostly aggrade the PSEs in PB and ERS, but PSEs appear to prograde where the shelf is strongly uplifted and eroded (WWS). Aggrading PSE geometries occur in unsampled areas of the outer shelf (WL, AP). On the slope, geometries show high-relief paleo-slopecanyons (PB, WL, RS, WS), and in the WS, deposition of the massive Crary Fan (WS) began. On the rise, seismic facies indicate higher energy depositional environments with paleo-channel-levee systems ﬁrst developing (PB, WL, RS, AP, WS). Where drilled (PB, WL, RS), lithologies from this period have glacial components indicating onshore glaciers. Backstripping calculations indicate normal shelf water-depths (PB, RS). Uniform prograding shelf (early and middle Miocene): In many regions, seismic sequences uniformly prograde the continental shelf edge, with an up-section increase in the dips of foreset-beds (glaciomarine deposits) and variable erosion of topset strata (with diamicton) (PB, WL, RS, AP, WS). Sea-level stratigraphic control began to shift to ice-dominated stratigraphic control, with documented cyclic shelf erosion by grounded ice sheets. Initial regional erosion and overdeepening of East Antarctic shelves commenced (PB, RS, WL). Slope geometries indicate canyon shifting and inﬁlling (PB, WL, RS) and fan growth (WS), with rise geometries showing the construction of large drift mounds (WWS, PB, WL, AP) and channel-levees (WL, WS, AP, PB). Abundant contourite deposits (PB, AP) with some turbidites (AP, WL) are documented. Glacial and interglacial sediment volumes decreased on the rise (PB, AP), but increased on the slope (WS). Local and focused prograding shelf (late Miocene to Pleistocene): A prominent regional unconformity occurs in the late Miocene to early Pliocene across all margin segments (PB (A, PP12); WL (U8); RS (U2); AP (BGMS) WS (W5)). The unconformity marks a circum-Antarctic change from areas of uniform PSE progradation to local-arcuate and broadly focused PSE progradation into small and/or overlapping upper-slope fans (e.g. QML, WL, AP) and broad trough-mouth fans (e.g. PB, WS) lying at or near the end of cross-shelf troughs. Strong regional shelf erosion, by both narrow and wide ice streams early in this period, was followed by regional deposition of topset banks composed principally of glacial diamicton (PB, AP, WL, WS?) and glacial marine (RS) sediment. Periods of pelagic sedimentation on the shelf indicate open water and sea ice. On the slope, foreset dips steepened as fans formed above the unconformity. On the rise, sedimentation rates decreased (PB, AP) and depocentres shifted progressively landward, moving from the rise to beneath the slope (PB, WL, AP,

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WS, AP?). Stratigraphic control was dominated by episodic grounded ice, with sediment deposition by ice and sediment distribution by ocean-currents. Sediment drape (late Pleistocene and Holocene): Thin well-layered acoustic units commonly inﬁll shelf depressions and drape across the slope and rise (All regions). These units provide a record of pelagic sedimentation from the last few interglacial and glacial periods. Deep inner-shelf basins, in particular, trap Holocene biogenic sediments with that were deposited at very high sedimentation rates; these biogenic sediments yield an ultra-highto high-resolution (i.e. decadal to millennial) record of climate variability (AP, WL, PB).

5.8.2. Long- and Short-Period Transitions in the Geologic Record The seismic stratigraphic record provides a regional framework that illustrates distinct changes in the morphology of the Antarctic margin over the last 60 m.y. Drill cores provide the direct ‘ground truth’ geologic record of both long-period (m.y.) and short-period (k.y.) transitions (Table D-2). Some of the transitions in drill cores are reﬂected in the seismic stratigraphic framework and others are not. All transitions, however, are important in deciphering Antarctic paleoenvironments. In this discussion, we focus on lithostratigraphic changes in drill cores, and leave the discussion of isotopic, biostratigraphic and other relevant variations to authors of other chapters of this book. Our intent is to use the proximal lithostratigraphic drilling record from the continental margin to independently evaluate paleoenvironmental history, where possible. Drill cores from two segments of the continental shelf provide a general lithostratigraphic framework for the long-period systematic transition from non-glacial (Mesozoic) to ﬂuctuating glacial and interglacial (late Cenozoic) paleo-depositional environments. At the mid-shelf of Prydz Bay, cores document the changes from subaerial non-glacial (Cretaceous) to ﬂuvial/ lagoonal early glacial (late Eocene) to shallow marine early glacial (early Oligocene) to subglacial deep shelf (early Pliocene) to interglacial openmarine (Holocene) conditions. A similar transition is documented in the Ross Sea (McMurdo Sound), from subaerial non-glacial (Mesozoic) to shallow marine early glacial (early Oligocene) to ﬂuctuating subglacial and marine glacial (early Miocene to Holocene) to interglacial open-marine (Holocene) environments. Large hiatuses exist in the shelf cores. Improved resolution of the continuity and timing of changes is seen in drill cores from the continental rise, where geologic continuity and core recovery

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Table D2-2: Lithostratigraphic transitions in the geologic records from drill cores from the antarctic margin (listed by duration and decreasing inferred age). Feature

Regions

Long-period changes (m.y.) On the shelf: up-section PB lithologic changes from alluvial to ﬂuvial to shallow marine to marine glacial to subglacial On the rise, sedimentation PB, AP, rates of hemipelagic WS sediments decrease smoothly while on the slope, sedimentation rates increase stepwise (i.e. in distinct stratigraphic units) On the rise, up-section PB increases in IRD, diatom content; shift kaolinite and glauconite Short-period changes (k.y.) On the inner shelf: cyclic RS changes from diamict (glacial) to glacial marine (interglacial) facies (in early Miocene at Milankovitch frequencies) On the rise: cyclic changes PB, AP from terrigenous (glacial) to biogenic (interglacial) facies at Milankovitch (PB) and similar variable (AP) frequency On the slope: shift in PB glacial sediment clast type: sandstone to granite a

The principal process is listed.

Processa Subsiding graben/ shelf with increasing ice to the shelf

Timing Cretaceous to Pliocene

Likely decrease in early Miocene to onshore sediment early Pliocene supply coincident with a shift in deposition from the rise to the slope

Erosion of shelf basins by grounded ice sheets

middle Miocene (B17–14 m.y.)

Glaciers ﬂuctuating early Oligocene to onto and off the middle Miocene shelf during glacial and interglacial times

Glaciers ﬂuctuating early Miocene to onto and off the early Pliocene shelf during glacial and interglacial times Likely change in ice Between 1.1 Ma and source area: 780 k.y. offshore to onshore?

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are greater than from the shelf or slope. On the rise, there is a distinct up-section change in seismic character from well layered below to channellevy development above (all areas) that is widely inferred due to a large inﬂux of sediment when onshore ice sheets initiated in late Eocene to early Oligocene time. Here, the pre-ice to glacial transition has not yet been sampled by drilling. Yet, higher in the stratigraphic section of the rise, drill cores show a clear long-term parabolic decrease in the sedimentation rates within sediment drift deposits from the early Miocene (PB: 10-fold decrease) and the late Miocene (AP: 6-fold decrease) to the present. The large decreases occurred when the PSEs were prograding over distances of several tens of kilometres and aggrading up to several hundred metres, although the detailed timing of the prograding and aggrading is unknown. Regardless, the distinct changes in geometries of seismic sequences beneath the outer paleoshelves in PB and AP are not seen as abrupt changes in sediment deposition rates on the rise. Drilling on the slope (WS) shows large incremental increases in sedimentation rates for this same general period (i.e. early Miocene to early Pliocene), indicating that sediment coming from the shelf may not have reached the rise. The decreases in sedimentation rates on the rise may also reﬂect a decrease in the amount of sediment being eroded from onshore and shelf areas. A notable long-term transition occurs in middle Miocene sediments (17–14 Ma) from the rise (PB). Up-section increases in IRD, diatom content, and recycled organic matter, along with changes in the types of clay and the ﬁrst appearance of glauconite, point to greater ice nearby and initial erosion of shelf sedimentary basins. Evidence for strong erosion on the shelf is also seen in truncated foreset strata beneath the outer PB paleo-shelf. RS shelf drill cores are marked by a long hiatus, from mid- to late-Miocene. The hiatus and truncated shelf reﬂectors point to an increase in shelf erosion and overdeepening in the RS, similar to the erosion recorded on the rise at PB. Short-period ﬂuctuations in shelf and rise drill cores provide the strongest evidence that erosion onshore and on the shelf by ﬂuctuating grounded ice sheets was the mechanism for sediment supply and distribution by glacial processes, as inferred from seismic-reﬂection data. On the RS shelf, cyclic ﬂuctuations in glacial diamict and interglacial glaciomarine lithofacies at Milankovitch frequencies are documented for lower Miocene nearshore facies at the front of the Transantarctic Mountains, and resulted from ice advancing onto and retreating off the shelf at this time (Naish et al., 2001). On the rise (PB and AP), alternating dark- and light-coloured lithofacies with varying amounts of terrigeneous (dark) and biogenic (light) components are described throughout the lower Miocene to lower Pliocene intervals from

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visual observation of cores, downhole logging and physical properties measurements (PB); similar compositions variations are described from upper Miocene to Pliocene intervals in the AP. The facies are inferred to be of glacial and interglacial origin, respectively, and occur at Milankovitch frequencies in PB and similar order-of-magnitude frequencies in AP. Hence, drilling on the rise in East Antarctica and in West Antarctica has provided similar geologic evidence for ﬂuctuating ice sheets on the shelves during the period of principal shelf progradation and aggradation – from the early Miocene to the early Pliocene. The geologic transition in the late Miocene to the early Pliocene is the initiation of broad and narrow shelf troughs and widespread banks, lobes and upper-slope fans along the Antarctic margin (all areas); this transition is difﬁcult to see in drill cores, but has been imaged seismically. On the rise (AP, PB), sedimentation rates decrease uniformly during this period and lithologies (e.g. clays, IRD, etc.) do not show systematic long-term changes. However, seismic geometries beneath the adjacent continental shelves show abrupt changes to rapidly prograding sections (S2/S3 beneath AP; PP-12 beneath PB). Elsewhere, drilling information is insufﬁcient to explain why large geomorphic changes on the shelf, probably due to changes in glacial regime, are not reﬂected in the rates or types of sediment delivered to the rise. 5.8.3. Sea-Level and Ice-Volume Changes Lithostratigraphic data from Antarctic margin drill cores show clear evidence on the shelf (PB, RS) for linked sea-level and ice-volume changes. This is best shown in the Oligocene through early Miocene record of cyclic glacial and interglacial lithologies near the coast in Cape Roberts cores (WRS) (Barrett, 2007). Lithostratigraphic data from the slope (PB) and rise (PB, AP) show additional direct evidence for cyclic ice-volume changes. Seismic-reﬂection data provide indirect evidence across the entire margin for the linked sea-level and ice-volume ﬂuctuations that have been noted by many investigators and have been modelled in the RS (Bartek et al., 1991) and presented conceptually for all margins (ten Brink et al., 1995). The Antarctic drill cores are too limited, however, to establish the timing, magnitude and extent of individual ice sheet advances onto the continental shelf, other than for the LGM. A comparison of the Antarctic margin proximal stratigraphic record with the global record of sea-level variations (and linked ice-volume variations) from coastal onlap, backstripping, and isotopic records since the middle

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Figure D-1: Graph showing sea-level and isotope curves and principal stratigraphic events for the Antarctic continental margin. Curves are from Miller et al. (2005). Events are from text and Tables D-1 and D-2. Curve A is for benthic foraminifera. Curve B is derived from stratigraphic backstripping (W9 M.a.) and isotopic measurements (0–9 M.a.); Curve C is from Haq et al. (1987). Curves B and C have different sea-level change scales. For some periods, Antarctic stratigraphic events correlate with isotopic shifts (e.g. PB: early Oligocene unconformity and ﬁrst ice sheets at the coast and midMiocene lithology changes and ice buildup) and with the Haq Curve (e.g. mid-to-late Miocene shelf erosion and early Pliocene and younger slope fan development with long-term sea-level lowerings).

Eocene is shown in Fig. D-1. Large differences appear between the global sea-level curves (see Miller et al. (2005) for the explanation), yet the Antarctic stratigraphic features can potentially be linked to parts of all of the curves principally because of the current uncertainty in ages of Antarctic features,

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especially for the Paleogene. The closest links of Antarctic features with the global curves are: Late Eocene and early Oligocene lowering of sea-level (ﬁrst glaciers into PB and RS), followed by sea-level rise in the early Oligocene (ﬂooding of PB). A long period from the Oligocene into the late Miocene of cyclic sea-level, seen in cyclic coastal deposits (WRS) and as glacial/interglacial rise-drift deposits (PB, AP) Abrupt sea-level lowering in the middle Miocene, seen in lithologies of rise-drift deposits (PB), followed by further sea-level lowering (with ice buildup) and enhanced erosional deepening and prograding of continental shelves by glaciers (as seen in seismic proﬁles from all areas). The systematic decrease in sea-level (and increase in ice) from the midMiocene to the present that corresponds with the decrease in sedimentation rates on the rise and increase in sedimentation rates on the slope. This is most pronounced since early Pliocene time, when an extensive system of shelf erosional troughs and upper-slope fans developed. Greater resolution of the link between ice-volume variations and sea-level changes requires further drilling on the Antarctic margin.

5.9. Summary The Antarctic continental margin holds a thick Cenozoic sedimentary section that is characterized by both long-period and short-period lithostratigraphic transitions, which are seen locally in drill cores and regionally in seismic-reﬂection data. Age resolution is inadequate to link individual stratigraphic events, but is sufﬁcient to make general statements about glacial history. The transitions point to the last 40 m.y. being a period of increasing glaciation and sediment distribution by glacial processes via short-period ﬂuctuations (e.g. Milankovitch frequencies) of grounded ice sheets across the continental shelf and accompanying sea-level changes. The proximal history is generally similar to that of distal proxy records from isotopic studies in adjacent ocean basins (e.g. Zachos et al., 2001) and from stratigraphic studies on low-latitude continental margins (e.g. Miller et al., 2005, 2000). Key inferences from the Antarctic margin for the Cenozoic, based on published data and inferences from extensive seismic-reﬂection and limited drilling records, are: Although tectonic histories differ around the Antarctic margin, similar geomorphic features (e.g. overdeepened and foredeepened seaﬂoor; broad

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erosional troughs, sediment fans and drift mounds) are seen everywhere on the margin, as a result of ubiquitous ﬂuctuating glaciers eroding and distributing sediments. East and West Antarctic margin segments may have similar glacial histories, based on similar geomorphologies and known ages offshore for glacial strata. Current differences result partly from lack of sufﬁcient drilling into likely Paleogene offshore sections beneath the margin. Seismic geometries and facies from all segments of the continental margin show evidence for up-section increases in the dynamic movement of sediment across the margin (e.g. shelf troughs, slope sediment fans, channel-levee systems) and along the margin (e.g. rise-drift deposits), reﬂecting increased glacial and ocean current activity from the Oligocene to the present. Stratal geometries of the continental shelf and slope were controlled principally by eustatic changes (with ice ﬂuctuations) from the Paleogene to about the middle Miocene, and thereafter principally by ﬂuctuating grounded glaciers (in tandem with sea-level changes) on the shelf, leading to the overdeepening and foredeepening of the shelf. Extensive prograding and aggrading of the continental shelf from the early Miocene to the latest Neogene is the principal result of sediment dispersal by ice sheets during glacial and interglacial periods at near-Milankovitch periodicities, as documented from drilling of drift deposits on the continental rise in East Antarctica (PB) and West Antarctica (AP), and from nearcoastal sequences (RS). The principal locus of sediment deposition on the margin has shifted from the outer rise (and beyond) during the Paleogene, to the inner rise and slope in the early Miocene to early Pliocene, and to the mid slope thereafter. The depocentre shift reﬂects the increases in glacial activity (and increases in ocean currents) and decrease in sediment being supplied due to erosion of onshore and shelf areas. Speciﬁc circum-Antarctic glacial events in the evolution of the margin include: ﬁrst glaciers at the coast and initiation of channel-levee systems on the rise and the Crary Fan (early Oligocene); ﬂuctuating glaciers, initial rapid progradation of the continental shelf, and initial growth of drift mounds and large levees on the rise (early Miocene); onshore ice buildup and initial overdeepening of the continental shelves (middle Miocene); dynamic ice movements and initial widespread development of cross-shelf troughs and upper-slope fans (early Pliocene); widespread deposition of biogenic interglacial sediment in deep inner-shelf troughs (Holocene). Additional advances in our understanding of Antarctica’s glacial history and the varied effects of ice sheets on the paleoceanographic and

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lithostratigraphic processes of the Antarctic continental margin can only be achieved through additional offshore deep stratigraphic drilling studies, such as the current IODP, ANDRILL and SHALDRIL projects.

ACKNOWLEDGEMENTS We thank many people who collaborated, by sharing data and ideas, on geoscience research projects under the umbrella of the highly successful ANTOSTRAT project for nearly 15 years. This synthesis, which reﬂects our views, would not have been possible without the efforts of these many investigators, most of whom continue their collaborative Antarctic studies, now under the successor ACE project. We appreciate the helpful review comments by Larry Krissek and Peter Barrett on this chapter. We thank John Anderson for compiling and providing the supplementary references to provide further historic perspective to community efforts partly summarized in this chapter.

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Foldout RS-1: ‘Type section’ from the Eastern basin (modiﬁed from Cooper et al., 1995). The section is a compilation of seismic lines BGR-7 and IFP-208. RSS-1 are inferred pre-ice deposits, RSS-2 and -3 are early-glacial glaciomarine deposits, RSS-4, -5 and -6 are glaciomarine deposits of the ice-sheet growth phase, and RSS-7 and -8 are glaciomarine deposits of the polar ice-sheet phase. See Fig. RS-1 for location.

Foldout RS-2: ‘Type section’ from the Northern basin (modiﬁed from Cooper et al., 1995). The stratigraphy and origin of sediments are similar to those for the ‘type section’ in the Eastern basin (Foldout RS-1). RSS-1 is a basin-ﬁll sequence, RSS-2 and -3 mostly aggrade, RSS-4, -5 and -6 mostly prograde and have with eroded topset beds, RSS-7 and -8 have aggrading geometries. See Fig. RS-1 for location.

50 km

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Foldout PB-1: Seismic sections from eastern Prydz Bay (Section B-Bu). In the bay, Cenozoic sediments overlie Cretaceous (Surface K) and older sequences resting on basement. The shelf edge and upper slope prograded seaward through time becoming steeper. The slope and rise overlie a thick post-rift section, the latest parts of which are turbidite (inferred) and contourite (drilled) sequences deposited since the onset of glaciation. Location shown in Fig. PB-1.

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Foldout PB-2: Seismic sections from western Prydz Bay (Section A-Au). The very thick post-rift section includes thick contourite (drilled) and turbidite (inferred) mounds with mudwaves in places. The upper slope comprises the Prydz Channel Fan and the shelf shows progradation but only limited topset thickness. The base of Prydz Channel Fan is Reﬂector A of Mizukoshi et al. (1986) and PP-12 of O’Brien et al. (2004). Location shown in Fig. PB-19.

Foldout WS-1: Seismic line NARE-8508 along the front of the Crary Trough Mouth Fan, showing fan architecture (modiﬁed from Kuvaas and Kristoffersen, 1991). Proﬁle location is in Fig. WS-1.

Foldout AP-1: Seismic line BAS878-19 across the Antarctic Peninsula Paciﬁc margin, showing the shelf basin, prograding wedge and continental rise sediments between the continental slope and DSDP Site 325. A buried, early Miocene sediment drift is observed beneath the central rise (Herna´ndez-Molina et al., 2004, 2006b), but this line lies between two of the large drifts that have developed since the start of the middle Miocene (Rebesco et al., 1996, 1997, 2002). Oceanic basement ages are from Larter et al. (1997). Vertical exaggeration at the seaﬂoor is 23:1. Inset in upper right shows an interpretation of major features. Inset in lower left shows line location.

Developments in Earth & Environmental Sciences, 8 F. Florindo and M. Siegert (Editors) r 2009 Elsevier B.V. All rights reserved DOI 10.1016/S1571-9197(08)00006-2

Chapter 6

Numerical Modelling of the Antarctic Ice Sheet Martin J. Siegert School of GeoSciences, University of Edinburgh, Grant Institute, West Mains Road, Edinburgh EH9 3JW, UK

ABSTRACT Studies on Antarctic climate evolution have beneﬁted increasingly over the last 20 years from numerical ice-sheet modelling. Such activities have led to the testing of geological hypotheses concerning past ice-sheet changes and a better understanding of macro-scale glaciological processes through space and time. Here, numerical ice-sheet models are reviewed, and their use in understanding former changes in Antarctica is outlined.

6.1. Introduction Analyses of glacial geological data allow hypotheses concerning Antarctic Ice Sheet history to be formed. Testing these hypotheses can be undertaken in two ways. First, additional geological data can be acquired and, second, numerical ice-sheet modelling experiments can be performed. There are several advantages in taking the latter option over the former. Ice-sheet modelling allows time-dependent assessments of ice-sheet form and ﬂow across the entire continent. Furthermore, model results can be analysed to determine where the ice sheet is most sensitive to changes and, in doing so, where the best sedimentary records of such changes may occur. Finally, the cost of undertaking numerical ice-sheet modelling is cheap relative to Antarctic data gathering. Hence, the use of ice-sheet modelling to test geologically based hypotheses has increased in recent years, and indeed is the Corresponding author. Tel.: +44(0)131 650 7543; Fax: +44(0)131 668 3184;

E-mail: [email protected] (M.J. Siegert).

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focus of the Antarctic Climate Evolution programme. Ice-sheet models have limitations, however, concerning the necessary simpliﬁcation of ice ﬂow processes, and incomplete model input data (such as bed topography and climate forcing). Researchers need to comprehend these limitations when interpreting results. In this chapter, generic numerical ice-sheet modelling is reviewed, and ways in which Antarctic models may be validated are assessed. The chapter provides necessary background information to later chapters that discuss results from ice-sheet modelling.

6.2. Ice-Sheet Processes 6.2.1. Flow of Ice Direct observations of glacier motion and laboratory studies of ice rheology (the ﬂow and deformation of ice) have identiﬁed three main mechanisms by which ice masses move. These are by internal deformation of the ice itself, and through two processes which take place at the base of the glacier; basal sliding and the deformation of water-saturated weak sediments (Fig. 6.1). Flow by internal ice deformation takes place in all ice masses, and generally accounts for motion of a few metres per year. However, basal motion occurs only where the bed is at the pressure melting point such that water is present. Where basal motion takes place, glaciers may move at tens to hundreds, and sometimes a few thousand, metres per year. The ﬂow of ice sheets can be organized in general terms as follows. At the centre of an ice sheet, the ﬂow speed is very low (of the order of metres per year), and controlled by internal deformation. A particle of ice on the ice-sheet surface will be buried by subsequent snowfall and so will possess a relatively signiﬁcant vertical velocity component downwards into the ice. The ﬂow of ice radiates from the ice divide, where there is no lateral ﬂow. Ice-sheet interiors are characterized by a series of divides that deﬁne the margins of ice drainage basins. Ice sheets are effectively ‘‘drained’’ by fast-ﬂowing rivers of ice, known as ice streams, transporting ice from the interior to the ice margin (Bennett, 2003). The velocity of ice streams is often several hundred metres per year. Ice streams ﬂow quickly because water at their bases causes a reduction in subglacial friction allowing them to effectively slide across the subglacial topography, with ice deformation contributing only a small amount to the total velocity. The transition between the slow-moving ice sheet and the fastﬂowing ice streams has been shown recently to occur in ‘‘tributaries’’ several hundred kilometres inland from the margin (Bamber et al., 2000).

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rface

ier su

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resting on bedrock

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(b) Cold--based glacier resting on bedrock

Glacier bed Bedrock

rface Glacier su Ice flow Sliding

on deformable sediment

Glacier bed Deforming horizon

Subglacial deformation

(c) Warm-based glacier resting

Stable bed

Figure 6.1: Processes controlling the ﬂow of large ice sheets. Ice will ﬂow through internal deformation in all situations. If the ice rests on bedrock and is warm based (a) then basal sliding can occur. If the base is frozen (b) then only ice deformation will take place. If the base is warm and loose unconsolidated sediments are present (c) then their own deformation can add to ice deformation and basal sliding as a contribution to ice ﬂow. Each of these processes can be accounted for in ice-sheet models. Adapted from Siegert (2001). 6.2.2. Mass Balance Antarctica gains mass through accumulation of snow in two ways. Direct precipitation of snow occurs close to the ice-sheet margin, where accumulation rates can reach B1 m per year. This is in contrast to accumulation rates in the interior, which are far lower (centimetres per year) and are associated

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with solid precipitation direct from water vapour. Broadly speaking, ice loss in Antarctica is through two main processes: iceberg calving and sub-iceshelf melting (Jacobs et al., 1992). Consequently, ice-shelf processes are critical to the overall mass balance of the Antarctic Ice Sheet. On grounded ice, surface sublimation, which forms blue ice zones, and basal melting, responsible for at least 145 subglacial lakes (Siegert et al., 2005a), are not thought to contribute signiﬁcantly to the overall ice-sheet mass balance.

6.2.3. Isostasy When a glacial load is placed on the Earth’s crust, the weight of ice will act to displace the crust into the viscous asthenosphere (beneath the lithosphere), which adjusts towards isostatic equilibrium according to Archimedes Principle, on time scales of a few thousand years. The process by which this is done involves ﬂow in the asthenosphere and elastic deﬂection of the lithosphere. When the ice load is removed (deglaciation), the asthenosphere and lithosphere will relax back to their original state (i.e. isostatic uplift or recovery). There are several ways in which glacial isostasy can be accounted for in ice-sheet models, and a thorough description of the problem is provided in Le Meur and Huybrechts (1996).

6.3. Ice-Sheet Models Numerical ice-sheet models have been used extensively since the mid-1980s to aid the reconstruction of former ice sheets, to comprehend existing ice sheets and to forecast future behaviour. The principle behind numerical icesheet modelling is that an ice sheet can be divided into a number of ‘‘ice columns’’. Each of these columns represents a ‘‘cell’’ in the model’s 2D horizontal ‘‘grid’’. Ice-sheet models are usually arranged in a ‘‘loop’’ that begins by applying a series of algorithms, determining the ﬂow of ice, mass balance and interaction with the Earth, in each cell. The loop is completed by a ﬁnal equation (widely known as a ‘‘continuity’’ equation), used on the full grid, to calculate the interaction and ﬂow of ice between cells. Each iteration of the loop advances the model through one time step. The accuracy of the model depends in part on the width of the grid cells, and the time step length. For continental-scale ice-sheet models, where the time-dependent change in ice-sheet behaviour needs to be calculated over several thousand years, the grid cell width is usually between ﬁve and twenty kilometres, and the loop

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time step is between 1 and 10 years. Because ice-sheet modelling is a valuable tool for palaeo-glaciologists, it is appropriate here to outline brieﬂy how icesheet models are organized. Most ice-sheet models are centred on the continuity equation for ice (Mahaffy, 1976), where the time-dependent change in ice thickness is associated with the speciﬁc net mass budget as follows: dH ¼ bs ðx; tÞ r F ðu; H Þ dt where F(u, H) is the net ﬂux of ice (m2 per year) (the ﬂux of ice being the product of ice velocity, u, and ice thickness, H), r the divergence operator and bs the speciﬁc mass budget term (involving surface mass balance, iceberg calving and ice-shelf basal melting where necessary). In the simplest models, the depth-averaged ice velocity, u (m s1), is calculated by the sum of depthaveraged internal ice deformation and basal motion (Fig. 6.1). Such models make the so-called ‘‘shallow-ice’’ approximation and assume that ice ﬂow can be described by a single horizontal vector. More complex models involving the full three-dimensional ﬂow of ice require a sophisticated solution to the stresses and strains which govern the ﬂow of ice (named ‘‘high-order stress models’’). Numerical ice-sheet models calculate glacial processes in a discrete manner (regardless of their sophistication). These processes are linked, however, such that feedback mechanisms can exist between the calculations. The relationship between distinct ice-sheet processes can be summarized by a ﬂow diagram (Fig. 6.2). A good example of a feedback mechanism involves glacial isostasy, which acts to adjust bedrock elevation due to ice loading and, hence, the ice-sheet surface elevation. Surface mass balance and ice velocity will be modiﬁed by surface elevation changes, which in turn feed back on ice thickness and, hence, ice loading. Another feedback loop acts to change thermal conditions at the base of ice sheets. This is important because the ice velocity due to rapid basal motion (e.g. sliding) is inﬂuenced by the thermal regime. In a cold-based ice-sheet, ice velocities are generally low and so, theoretically, the ice sheet is allowed to build up with a minimum of basal sliding. However, as the ice sheet thickens, the basal temperature is likely to increase (because both surface temperatures and vertical temperature gradients remain relatively constant). Once the base of the ice sheet starts to melt, rapid basal motion becomes possible. Fast ice ﬂow leads to more ice being advected to the ice-sheet margins and, so to ice-sheet thinning. Thinner ice then causes a reduction in the temperature of the icesheet base and possibly to the curtailment of rapid basal motion. This feedback process is complicated because the thermal regime is inﬂuenced by

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Basal marine melt

Initial precipitation

Precipitation

Isostasy Initial bedrock

Net mass balance

Ice thickness

Calving and maritime melt

Bedrock elevation Eustasy

Ablation and effectiveness

Latitude

Ice velocity

Surface slope

Ice surface elevation

Imposed temperature change

Figure 6.2: Flow diagram of the operation of an ice-sheet model. Adapted from Siegert (2001). the basal heat gradient, which is in part related to the heat due to basal sliding. Thus, rapid basal motion may maintain warm-based conditions even if the ice sheet thins. One might consider when it is appropriate to employ simple ﬂow models versus complex models. Both have a role to play in understanding the behaviour of large ice sheets. The beneﬁt of simple models, such as those using the ‘‘shallow-ice approximation’’, is twofold. First, in central regions of the ice sheet, the approximation works well at a horizontal scale greater than ﬁve ice thicknesses. In other words, there is little change in output at this scale between a simple and a complex model. Second, they can be run quickly, which means multiple runs and sensitivity experiments can be undertaken with minimum computing costs. More complex models are required to solve particular ice ﬂow issues, such as the dynamic change between ﬂoating and grounded ice that occurs at the ice-sheet/shelf transition (Schoof, 2007) and over subglacial lakes (e.g. Pattyn, 2003). Given that recent changes to the Antarctic Ice Sheet have been observed at the marine ice-sheet margin (for example, in the recent IPCC report in 2007), understanding ice-sheet response to future warming scenarios is likely to require complex models. High-order models are also able to run at

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greater spatial accuracy, meaning that important features such as ice streams are resolved. However, given the uncertainty in model inputs regarding palaeo-ice sheets (see below), it is arguable whether the complexity accounted for in high-order stress models adds any value to the output of simpler models.

6.4. Model Inputs In numerical modelling (ice sheet, ocean or atmospheric), the quality of model output is dependent on the quality of model input. The most important data used as boundary and forcing conditions in an ice-sheet model are subglacial topography and surface mass balance. For present-day ice sheets, the quality of these model inputs is dependent on ﬁeld data which, for some areas, are noticeably absent. Airborne radar surveying of large ice sheets provides the only viable method of acquiring information on the icesheet base at a continental scale. Such surveys are organized as a series of gridded ﬂightlines. Although the technique is sound, there are currently two limitations to datasets collected by radar surveying. The ﬁrst is that information between ﬂightlines is absent. Interpolation software is used to infer the topography in such regions. The second limitation, which is acute in Antarctica, is that large portions of some ice sheets remain to be surveyed. Thus, subglacial topography in these regions has to be estimated, often from just a few seismic measurements, and is subject to a high level of potential error. To date, Antarctic Ice Sheet models have been run over a topographic DEM based on an incomplete coverage of bed elevation data. There have been several recent advances to alleviate this problem. Since the 1970s, when radar data from over 400,000 km worth of ﬂightlines were collected over a wide area of Antarctica (albeit at a coarse line spacing), there have been several surveys over much smaller regions with close ﬂightline spacing. The most recent depiction of Antarctic topography was assembled by the BEDMAP consortium (Lythe et al., 2000). While this database quantiﬁes topography well in many places, it remains restricted in two ways. First, there are several large ‘‘data gaps’’ where very little is known about sub-ice topography. Second, even in regions where coverage is good, radar transects are often separated by several kilometres (often tens of kilometres) across which interpolation of data remains necessary. Comparison of the topography measured by recently acquired radar data with the interpolated information from BEDMAP reveals that, in some places, the errors in BEDMAP are up to the order of hundreds of metres (Welch and Jacobel, 2003).

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Mass balance of large ice masses is also very difﬁcult to establish accurately over a wide area. Although it is possible to measure annual mass balance at speciﬁc sites, there is no quick and easy method by which large-scale data can be obtained. Satellite observations, used in conjunction with direct measurements, provide a means of estimating mass balance across a wide area, and have been used to assess Antarctic mass balance (Vaughan et al., 1999; Arthern et al., 2006). Such analysis only provides an estimate of the current situation, however, which is of less use when trying to model past or future scenarios. As this is a real problem for Antarctic Ice Sheet models, many are coupled with models of surface mass balance. The accuracy of ice-sheet model results will then be dependent on the quality of the surface balance results. Similar problems exist when accounting for the processes of iceberg calving and ice-shelf basal melting in ice-sheet models.

6.5. EISMINT The majority of ice-sheet models have been developed by individual researchers or within small institution-based groups. This ‘‘do-it-yourself’’ history is in contrast to that of large community-based models, which dominate research in the ﬁelds of meteorology and oceanography. The resulting diversity of ice-sheet models is both a strength, in that speciﬁc models were honed to speciﬁc applications, and a weakness, in that no consensus on the basic behaviour of the models existed. The European Ice Sheet Modelling Initiative (EISMINT) was developed to address this diversity. Siegert and Payne (2001) provide details of how the EISMINT programme has assisted the development of ice sheet models. The following section is summarised from Siegert and Payne (2001) to illustrate the purpose of EISMINT and its achievements. EISMINT had three main aims: ﬁrst, to deﬁne and perform intercomparison exercises to help establish consensus on basic model prediction; second, to identify good practice in ice-sheet modelling; third, to develop the next generation of ice-sheet models. The EISMINT project was funded by the European Science Foundation in two phases, which ran from January 1993 to December 1995 and from January 1996 to December 1997. The funding allowed four major international meetings to concentrate on model intercomparison. It also allowed a series of smaller meetings on basal processes, ice rheology, ice–climate interactions, ice–oceanic interactions, ice–lithosphere interactions and former ice sheets.

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The aims of model intercomparisons were threefold. The ﬁrst was to test the effects of numerical implementation on model prediction. This required the physics incorporated into models to be tightly constrained, along with the values of model parameters and boundary conditions. A wide variety of numerical techniques has been employed by ice-sheet modellers. For example, several ﬁnite-element models existed, although the majority of ice-sheet models employed ﬁnite difference schemes. Within the latter class, a variety of methods were available to deal with the basic nonlinearity of ice ﬂow. The EISMINT intercomparisons explained, for the ﬁrst time, the effect of the type of numerical solution on model prediction. The second was to determine the effect which the many poorly constrained physical parameters had on the overall prediction of the models. The third was to model ice masses through a given time-dependent climate change scenario. Examples included the response of ice sheets to stepped changes in forcing, glacial– interglacial growth and decay, and the response to future, anthropogenic change. Particular attention was paid to how models replicated and forecasted Antarctic Ice Sheet dynamics. 6.5.1. EISMINT Level One and Level Two Ice-sheet model intercomparison was designed to be as inclusive as possible, in order that results from models with one and two horizontal dimensions, which may or may not have included internal-temperature evolution, could be evaluated (Huybrechts et al., 1996, and papers within Annals of Glaciology, Vol. 23). Model boundary conditions and parameter values were prescribed as much as possible so that observed differences in output could be interpreted purely in terms of the numerics of the model. The intercomparison exercise involved several iterations in which model results were submitted for analysis. A consensus set of results gradually emerged from this process. Many of the differences identiﬁed initially between models arose for non-scientiﬁc reasons (such as ambiguity in the experiment descriptions). One problem encountered was that results tended to converge towards median values as outlying models were modiﬁed. The modiﬁcation process may have been genuine (for instance, in ﬁnding coding errors) but may also have been driven by a perceived need to conform with the bulk of models regardless of whether they were correct. In this way, genuine differences could have become obscured. This problem was offset somewhat by recognizing analytical solutions to ice ﬂow equations, taken as ‘‘truth’’. While the number of analytical solutions available is quite limited for ice-sheet models, their use was an important feature of the intercomparison process.

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In general, most models performed similarly in terms of their predictions of ice-thickness evolution, internal-temperature evolution and ice velocity patterns, but some differences were observed. Interpretation of these differences showed the effect of numerical solution on model predictions. Two groups of models were identiﬁed: those which employed a more stable scheme to solve the continuity equation were found to be less accurate in comparison to a more precise formulation which was, however, less stable numerically. In this way, the EISMINT programme was able to establish ﬁrm benchmarks for model types and their performance. 6.5.2. EISMINT Level Three Simulations of the Antarctic Ice Sheet took place in the third level of the EISMINT exercise during 1997. The aim of this level was to compare model predictions of the past and future evolution of these ice masses with the minimum amount of constraint on the details of the individual models. Three experiments were undertaken. The ﬁrst compared results from simulations of the present-day ice sheet. The second compared results from glacial–interglacial experiments. The third analysed differences between forecasts of future icesheet behaviour (over the next 500 years) (see http://homepages.vub.ac.be/ Bphuybrec/eismint.html for details). Results for the ﬁrst experiment were encouraging, as all four models revealed similar output. However, major differences were observed in the second and third experiments. Speciﬁcally, ice-sheet volumes predicted for both the Eemian interglacial (started at ca. 125 ka and ended at ca. 110 ka) and Last Glacial Maximum were noticeably different. Although only two of the four models yielded similar results, they were recommended as a benchmark for subsequent studies, presumably because other models had fatal ﬂaws that became apparent in the experiment, allowing a level of ‘‘veriﬁcation’’ for Antarctic Ice Sheet models. 6.5.3. Outcomes from EISMINT There is no doubt that the EISMINT programme has led to agreement and commonality regarding ice-sheet modelling. One important outcome is the establishment of a community within the ﬁeld of glaciology, which can develop ice-sheet modelling in future. Importantly, the EISMINT programme has led to the availability of free software, in the form of the GLIMMER model (http://forge.nesc.ac.uk/projects/glimmer/), based on the

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‘‘best practice’’ elements established through model intercomparison exercises. The full development, documentation and release of GLIMMER was a consequence of the GENIE Earth System Model project (www. genie.ac.uk), which used an ice-sheet model built on the validated procedures established by EISMINT. It is very likely that this model will lead to an increase in ice-sheet modelling activity and, thus, its use in understanding past changes in Antarctica. The next phase of ice-sheet model intercomparisons has recently been established in the ice sheet Model Intercomparison Programme (ISMIP). In ISMIP, the performance of next generation of icesheet models (accounting for the full three-dimensional ﬂow of ice) will be analysed. It will also help to establish community agreement on how icesheet models predict the cryosphere to respond to future climate change scenarios (details of ISMIP can be found at the following website: homepages.vub.ac.be/Bphuybrec/ismip.html).

6.6. Comparing Ice-Sheet Models with Antarctic Glaciological Data Despite most ice-sheet models being based on well-established physical assumptions about the ﬂow of ice, these assumptions, and the data used as input, may oversimplify the actual glaciological situation. The result is mismatch between model results and real glaciological measurements. When comparing model output with glaciological data, the scale of the datasets needs to be similar to the scale of the ice-sheet model output (e.g. averaged over 5–20 km) for the exercise to be meaningful. Real data at a ﬁner resolution than this show how ice-sheet models simplify the actual ﬂow of ice sheets. Determining the difference between model output and these real data represents an important way in which to assess the validity of ice-sheet models. As ice-sheet models become more sophisticated, it will be necessary to test their output against a variety of ice-sheet measurements for validation purposes. This section outlines datasets currently available in Antarctica that have yet to be compared fully with ice-sheet model output, but should be used for the future validation of ice-sheet models (Siegert and Payne, 2001). 6.6.1. Surface and Sub-Ice Morphology, and the Flow of Ice in Central Regions The advent of satellite altimetry has resulted in the determination of the surface morphology of ice sheets to a high degree of accuracy. At present,

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most ice-sheet models of Antarctica work at a 5–20 km scale. These models replicate the broad shape of the ice sheet well. However, real morphology at this or a ﬁner scale cannot be modelled at present. ERS-1 altimetry of the Antarctic Ice Sheets reveals that there is a complex morphology on the ice surface related to the ﬂow of ice. A good example of where ice-sheet model results (Huybrechts, 1990) match well with large-scale ice-sheet features, yet poorly with sub-grid cell morphology, is at Dome C, East Antarctica (Fig. 6.3). The ice surface around Dome C has a generalized surface slope of about 0.081. However, there are a number of regions where the slope reduces to less than 0.011. These ‘‘ﬂat’’ surfaces are caused by ice ﬂow over subglacial trenches and/or subglacial lakes (e.g. Lake Vostok). In either case, the ﬂow of ice is altered from the base-parallel shearing that most ice-sheet models account for, by longitudinal extension. In order to solve this problem, an accurate representation of subglacial topography is required. Although there is a

Figure 6.3: The location of Antarctic subglacial lakes. Adapted from Siegert et al. (2005a).

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major programme aimed at establishing an up to date bedrock elevation for Antarctica (e.g. BEDMAP), the resolution of these new data is not ﬁne enough to solve the problem of complex ﬂow of ice over subglacial topography highlighted here. 6.6.2. Radar Layering and the Internal Flow of Ice Some ice-sheet models are capable of calculating the ﬂow of ice in three dimensions. Ice penetrating radar information can be used to verify calculated ﬂow paths because it reveals internal layering which is assumed to be isochronous. Internal layering is detected by changes in the electrical properties of ice. However, there are three different ways by which these changes can occur, yielding three main types of internal layering. The ﬁrst type of internal layer is when there are changes in the density of ice. This is a dominant process at ice depths less than 700 m. However, below this level, the density of ice does not change very much and so internal reﬂections must be caused by other processes. The second form of layering is caused by the acidity of ice. Layers with an acidity in excess of the normal background level of glacier ice are formed when the aerosol products of ancient volcanic activity are incorporated in the snow chemistry on the former ice surface. This acid snow is subsequently buried by later snow fall to its present-day position, several tens of thousand years later. These acid layers are therefore isochronous surfaces. The third type of layering is where there are changes in the crystallography of ice. Such layering is thought to develop from acidic layers in the presence of enhanced stress across the stoss face of subglacial hills (Fujita et al., 1999). Internal layers are often continuous across large sections of the Antarctic Ice Sheet (Siegert, 1999). Because they are isochronous, their patterns can be used to match the 3D ﬂow of ice calculated in numerical models. However, as yet, very few ice-sheet modelling investigations have utilized this potentially powerful measurement. Leysinger Vieli et al. (2007) have recently identiﬁed a means by which internal layering can assist large-scale ice-sheet modelling, and this marks an important new area of model validation. 6.6.3. Location of Subglacial Lakes and Ice-Sheet Thermal Conditions Subglacial lakes are evidence for melting at the ice-sheet base. They can, therefore, be used to validate the thermal conditions calculated by ice-sheet models. Over 145 subglacial lakes have been identiﬁed at the base of the

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13 million km2 Antarctic Ice Sheet (Siegert et al., 2005a). The bulk of the lakes are located in the interior of the Antarctic Ice Sheet, at or close to ice divides in the form of domes and ridges, where ice velocity is low (Dowdeswell and Siegert, 1999). Almost 60 per cent of lakes are found within 200 km of an ice crest, remembering that ice ﬂowlines from divide to coastal margin are often over 1,000 km long in Antarctica (Fig. 6.3). Only about 15 per cent of subglacial lakes are positioned more than 500 km from an ice divide. At least 16 subglacial lakes occur at locations which are near the onset of enhanced ice ﬂow, some hundreds of kilometres from the ice-sheet crest. An example is provided by three subglacial lakes near the onset of fast ﬂow into Byrd Glacier. Byrd Glacier is fast ﬂowing and drains a very large interior ice-sheet drainage basin into the Ross Ice Shelf. These subglacial lakes are similar in size and depth to the small and probably shallow lakes found in major subglacial basins in the ice-sheet interior. Siegert and Bamber (2000) recognized subglacial lakes at the heads of ice stream tributaries (some of which are several hundred kilometres in the ice-sheet interior) as being evidence of warm-based conditions along the entire length of the enhanced ﬂow unit. Huybrechts (1990) used the location of subglacial lakes, as evidence of warm-based conditions, to verify the broad-scale thermal character of the icesheet base across East Antarctica. Basal melting conditions were predicted across both the centre of East Antarctica and at the base of ice streams. However, because of the smoothed bedrock topography used as model input, individual lakes, or the outline of speciﬁc lake regions, were not able to be matched well. There is, therefore, plenty of scope to better use subglacial lakes in the future veriﬁcation of ice-sheet model thermal conditions. Further validation of the ice-sheet thermal regime can be made by comparing results with temperature proﬁles from ice cores. However, this comparison is limited by the small number of deep ice cores and the lack of penetration to the lower layers of the ice sheet in most cases. 6.6.4. Satellite Interferometry One aspect of ice-sheet modelling that is in need of independent datasets for comparison and validation is the calculated velocity of ice. Until recently, such data have been scarce over large ice sheets. However, recent interferometric Synthetic Aperture Radar (SAR) InSAR techniques have been applied to glaciers and ice sheets to reveal the surface velocity continuously across wide areas. Work to date has shown that development of ice streams in West Antarctica seen through InSAR data (Joughin et al., 1999) compares well with

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calculations of Antarctic ice ﬂux (Bamber et al., 2000). Maps of surface ice velocities are currently being assembled for the whole of Antarctica, which may be a vital dataset for model veriﬁcation in future. InSAR data revealing the surface velocity of ice sheets also have an application in modelling subglacial sliding. In areas where the internal icesheet deformation is known to occur in a simple manner, the difference between the InSAR velocity and velocity due to internal deformation of ice will be equivalent to the component due to basal processes. Currently, models are incapable of accurately predicting subglacial sliding and, so, InSAR data may be of future use in establishing reliable algorithms for this process.

6.7. Ice-Sheet Reconstructions Once ice-sheet models have been veriﬁed and validated against the modern glaciological regime, they can be used to test geological hypotheses regarding past changes in Antarctica. Whereas the boundary conditions of the present ice sheet can be established reasonably well, past conditions are difﬁcult to quantify. This difﬁculty obviously increases with past time. Researchers should be aware, therefore, of the limitations of ice-sheet models, and the reliability of information they are able to provide, given necessary assumptions about the past environment. Despite this problem, ice-sheet modelling has been used in innovative ways to help understand past ice-sheet changes in Antarctica. Two examples of such activities are provided below.

6.7.1. Pliocene Ice Sheets Huybrechts (1993) provided the ﬁrst detailed modelling investigation into Pliocene Antarctic Ice Sheets. He tackled the problem of whether the East Antarctic Ice Sheet was stable at this time, or susceptible to signiﬁcant changes (both hypotheses have been developed from the geological record; see Miller and Mabin, 1998). Running a standard ice-sheet model, using the shallow-ice approximation under a selection of climate forcing parameters, Huybrechts showed that the East Antarctic Ice Sheet could withstand a mean annual temperature rise of around 101C and that several degrees in excess of this amount of warming was required to signiﬁcantly alter the ice-sheet conﬁguration (Fig. 6.4). When the ice sheet was forced to change, it ﬁrst did so predominantly across the Dome C region, where topography is lowest in East Antarctica. While the model was not used to understand the form and

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Figure 6.4: Numerical modelling results of the Antarctic Ice Sheet under a variety of climate warming scenarios. Taken from Huybrechts (1993) with permission from The University of Edinburgh. ﬂow of the Pliocene ice sheet, it did demonstrate that even with substantial climate warming, a large semi-continental-scale ice mass was predicted. Air temperatures needed to rise by around 151C of modern values in order to reduce the ice sheet to a series of small isolated ice caps. As such warming is unlikely for the Pliocene, Huybrechts (1993) concluded that the idea of widespread decay of ice at this time is not supported by ice-sheet modelling. Huybrechts’ results are relatively easy to replicate within a simple ice-sheet model. Siegert et al. (2005b) ran a basic shallow-ice approximation EISMINT model with a simplistic climate control. By raising the equilibrium

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line altitude (ELA) to 400 m above sea level (i.e. a climate warming scenario of over 151C), the East Antarctic Ice Sheet reduced in size across the Dome C region. Further increases in the ELA resulted in the complete decay of ice at Dome C, but retention of a large ice sheet on the higher ground focused at Dome A and Ridge B.

6.7.2. Pleistocene Ice Sheets There have been several models of recent (last glacial cycle) ice-sheet behaviour (e.g. Ritz et al., 2001; Huybrechts, 2002). Huybrechts’ model demonstrated the differences in behaviour between East and West Antarctica over the last 200,000 years, and showed these differences to be controlled to a large degree by topography (Fig. 6.5). The West Antarctic Ice Sheet is known as a marine-based ice sheet, with a bed that is well below sea level in most places (mean bed elevation B400 m). This is in contrast to the terrestrial ice sheet in East Antarctica, where the bed elevation is generally above sea level (mean bed elevation Bþ50 m). Marine ice sheets are thought to be susceptible to greater change than terrestrial ice sheets as a consequence of (1) grounding line retreats caused by the disintegration of ice shelves, (2) the deepening of the bed inland of ice-sheet margins, which can lead to a positive feedback of ice decay, (3) the level to which the ice sheet is held buoyant by the surrounding sea, which may be enhanced if sea level rises, and (4) the nature and stability of ice streams, which may be underlain by weak marine sediments leading to enhanced ﬂow speeds and complex dynamics. In the Pleistocene, the lower topography in West Antarctica was associated with signiﬁcant ﬂuctuations in ice-sheet size in contrast with a relatively stable East Antarctic Ice Sheet. While the ice stream conﬁgurations in East Antarctica remain largely similar (although they may reduce in ﬂux), signiﬁcant adjustments to the ice ﬂow regime are predicted in West Antarctica. Consequently, ice ﬂow in East Antarctica is associated strongly with topography, in which ice streams are most often constrained in topographic channels. In West Antarctica, the association between ﬂow and topography is less clear. Here, ice streams ﬂow over relatively ﬂat terrain where weak water-saturated sediments dominate. Lateral migration of ice stream margins, without topographic controls, is entirely feasible in such places. The outcome is that the spatial pattern of glacial erosion and deposition in East Antarctica appears to have been relatively consistent over the past few glacial cycles, and indeed back through several million years, whereas substantial changes are likely to have occurred in West Antarctica.

Figure 6.5: Numerical modelling of Antarctic ice surface elevation during the last glacial–interglacial, including the ice-sheet conﬁguration at the last glacial maximum. Taken from Huybrechts (2002) with permission.

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Although Huybrechts’ (2002) investigation represents the most advanced numerical analysis of Pleistocene ice sheets in Antarctica, the model used has several limitations that prevent it from resolving ice streams well, or ice shelves at all. Increasing the complexity of the model to solve this issue might not lead to improved results, however, due to greater uncertainties that remain in terms of bed conﬁguration and accumulation rates. The lesson to be learned is that glaciologists have an important choice concerning model complexity and the robustness of the model’s outputs, which must be made through consideration of the aims of the investigation. A broad-scale reconstruction at a continental level may still be best suited to a simple model, whereas information concerning time-dependent changes in discrete ice-sheet processes might be better served through a high-order model.

6.8. Summary Ice-sheet modelling allows quantitative predictions of how large ice masses behave and respond to environmental change. As many different ice-sheet models were developed in the early 1990s, the EISMINT programme was set up to establish best practice and commonality within the ﬁeld. Through intercomparison exercises, the EISMINT programme established appropriate controls on ice-sheet modelling. One of the results is the development of freely available ice-sheet modelling software, which can be used in future to help test geological hypotheses concerning past changes in Antarctica. Although the large-scale behaviour of ice sheets can be reproduced well by numerical models, there are limitations to the accuracy of numerical ice-sheet modelling at a ﬁner scale. This is because (1) ice-sheet models are derived from assumptions about the ﬂow of ice which may not be accurate at all places in the ice sheet and (2) subglacial topography and other model inputs may be insufﬁciently well known to allow the model to work at a ﬁne scale. The latter problem can be solved by obtaining high-resolution geophysical datasets of existing ice sheets to establish the subglacial topography, internal structure and surface elevation of the ice sheet. Such data acquisition has been ongoing for several decades. Ice-sheet models have been used to comprehend past changes, and test hypotheses based on geological data. For example, Huybrechts (1993) showed that climate warming of over 151C was required to decay the East Antarctic Ice Sheet into small discrete ice masses. As a result, the idea of restricted ice cover during the Pliocene was difﬁcult to support by ice-sheet modelling. In additional exercises, Huybrechts (2002) revealed the sensitivity

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of the East and West Antarctic Ice Sheets to climate changes was noticeably different. In East Antarctica, the ice sheet appears resistant to even quite large changes, whereas modiﬁcation to the size and shape of the West Antarctic Ice Sheet can be predicted with only small adjustments to boundary conditions. This sensitivity may characterize Antarctic glacial history throughout the Pleistocene and possibly much earlier. Ice-sheet models are used to predict how the cryosphere responds to and affects climate change. From climate information gathered from ice cores, there is reason to believe that ice sheets, oceans and the atmosphere interact with each other. It is therefore important that ice-sheet models be coupled with models of atmospheric and ocean circulation in order to predict how the ice–ocean–atmosphere system behaves in the future and has behaved in the past (e.g. DeConto and Pollard, 2003). Examples of such coupling are provided in later chapters, and these activities mark a clear way for future reconstructions of past ice sheets in Antarctica.

REFERENCES Arthern, R. J., Winebrenner, D. P., & Vaughan, D. G. (2006). Antarctic snow accumulation mapped using polarization of 4.3-cm wavelength microwave emission. J. Geophys. Res., 111, D06107. Bamber, J. L., Vaughan, D. G., & Joughin, I. (2000). Widespread complex ﬂow in the interior of the Antarctic Ice Sheet. Science, 287, 1248–1250. Bennett, M. R. (2003). Ice streams as the arteries of an ice sheet: Their mechanics, stability and signiﬁcance. Earth Sci. Rev., 61, 309–339. DeConto, R. M., & Pollard, D. (2003). Rapid Cenozoic glaciation of Antarctica induced by declining atmospheric CO2. Nature, 421, 245–249. Dowdeswell, J. A., & Siegert, M. J. (1999). The dimensions and topographic setting of Antarctic subglacial lakes and implications for large-scale water storage beneath continental ice sheets. Geol. Soc. Am. Bull., 111, 254–263. Fujita, S., Maeno, H., Uratsuka, S., Furukawa, T., Mae, S., Fujii, Y., & Watanabe, O. (1999). Nature of radio-echo layering in the Antarctic Ice Sheet detected by a two-frequency experiment. J. Geophys. Res., 104(B6), 13013–13024. Huybrechts, P. (1990). A 3-D model for the Antarctic Ice Sheet: A sensitivity study on the glacial interglacial contrast. Climate Dynamics, 5, 79–92. Huybrechts, P. (1993). Glaciological modelling of the Late Cenozoic East Antarctic Ice Sheet: Stability or dynamism? Geograﬁska Annaler, 75, 221–238. Huybrechts, P. (2002). Sea-level changes at the LGM from ice – Dynamic reconstructions of the Greenland and Antarctic Ice Sheets during the glacial cycles. Quaternary Sci. Rev., 21, 203–231. Huybrechts, P., Payne, A., & The EISMINT Intercomparison Group (1996). The EISMINT benchmarks for testing ice-sheet models. Ann. Glaciol., 23, 1–12.

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IPCC (2007). Climate Change 2007: The Physical Science Basis. Contribution of Working Group I to the Fourth Assessment Report of the Inter governmental Panel on Climate Change. In: S. Solomon, D. Qin, M. Manning, Z. Chen, M. Marquis, K. B. Averyt, M. Tignor, & H. L. Miller (Eds). Cambridge University Press, Cambridge, UK, 996 pp. Jacobs, S. S., Hellmer, H. H., Doake, C. S. M., Jenkins, A., & Frolich, R. M. (1992). Melting of ice shelves and the mass balance of Antarctica. J. Glaciol., 38, 375–387. Joughin, I., Gray, L., Bindschadler, R., Price, S., Morse, D., Hulbe, C., Mattar, K., & Werner, C. (1999). Tributaries of West Antarctic ice streams revealed by RADARSAT interferometry. Science, 286, 283–286. Le Meur, E., & Huybrechts, P. (1996). A comparison of different ways of dealing with isostasy: Examples from modelling the Antarctic Ice Sheet during the last glacial cycle. Ann. Glaciol., 23, 309–317. Leysinger Vieli, G. J. M. C., Hindmarsh, R. C. A., & Siegert, M. J. (2007). Three dimensional ﬂow inﬂuences on radar layer stratigraphy. Ann. Glaciol., 46, 22–28. Lythe, M. B., Vaughan, D. G., & The BEDMAP Consortium (2000). BEDMAP – Bed Topography of the Antarctic. 1:10,000,000 Map. BAS (Misc.) 9, Cambridge. Mahaffy, M. W. (1976). A three dimensional numerical model of ice sheets: Tests on the Barnes Ice Cap, northwest territories. J. Geophys. Res., 81, 1059–1066. Miller, M. F., & Mabin, M. C. G. (1998). Antarctic Neogene landscapes – In the refrigerator or in the deep freeze? GSA Today, 8(4), 1–2. Pattyn, F. (2003). A new three-dimensional higher-order thermomechanical ice-sheet model: Basic sensitivity, ice-stream development and ice ﬂow across subglacial lakes. J. Geophys. Res., 108(B8), 2382, doi:10.1029/2002JB002329. Ritz, C., Rommelaere, V., & Dumas, C. (2001). Modeling the evolution of the Antarctic Ice Sheet over the last 420,000 years: Implications for altitude changes in the Vostok region. J. Geophys. Res., 106, 31943–31964. Schoof, C. (2007). Ice sheet grounding line dynamics: Steady states, stability and hysteresis. J. Geophys. Res., 112, F03S28, doi:10.1029/2006F000664. Siegert, M. J. (1999). On the origin, nature and uses of Antarctic Ice-Sheet radioecho layering. Prog. Phys. Geogr., 23, 159–179. Siegert, M. J. (2001). Ice Sheets and Late Quaternary Environmental Change. John Wiley and Sons Ltd., Chichester, UK, 231 pp. Siegert, M. J., & Bamber, J. L. (2000). Subglacial water at the heads of Antarctic icestream tributaries. J. Glaciol., 46, 702–703. Siegert, M. J. & Payne, T. (2001). Validation of ice sheet models. In: Anderson, M. G. & Bater, P. D. (Eds). Model validation in hydrological science. John Wiley and Sons Ltd. Chichester, England, pp. 439–460. Siegert, M. J., Carter, S., Tabacco, I., Popov, S., & Blankenship, D. (2005a). A revised inventory of Antarctic subglacial lakes. Antarct. Sci., 17, 453–460.

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Siegert, M. J., Taylor, J., & Payne, A. J. (2005b). Spectral roughness of subglacial topography and implications for former ice-sheet dynamics in East Antarctica. Global Planet. Change, 45, 249–263. Vaughan, D. G., Bamber, J. L., Giovinetto, M., Russell, J., & Cooper, AP. R. (1999). Reassessment of net surface mass balance in Antarctica. J. Climate, 12, 933–946. Welch, B. C., & Jacobel, R. W. (2003). Analysis of deep-penetrating radar surveys of West Antarctica, US-ITASE 2001. Geophys. Res. Lett., 30(8), 1444, doi:10.1029/ 2003/GL017210.

Developments in Earth & Environmental Sciences, 8 F. Florindo and M. Siegert (Editors) r 2009 Elsevier B.V. All rights reserved DOI 10.1016/S1571-9197(08)00007-4

Chapter 7

The Antarctic Continent in Gondwanaland: A Tectonic Review and Potential Research Targets for Future Investigations F. M. Talarico1, and G. Kleinschmidt2 1

Dip. Scienze della Terra, Universita` di Siena, Via del Laterino 8, I-53100 Siena, Italy Institut f. Geowissenschaften d. Univesita¨t Frankfurt, Altenho¨ferallee 1, D-60438 Frankfurt/Main, Germany

2

ABSTRACT The geological record of the Antarctic continent has, for a long time, had a key role in supercontinent reconstructions. More recently, because of the welldocumented relevance of the polar regions’ processes in inﬂuencing the global changes of both ocean circulation and climate patterns, Antarctica has increasingly been of similar importance in the context of palaeoenvironmental and palaeoclimatic investigations, particularly those focused on the Cenozoic glacial evolution. After a description of the present-day geotectonic setting and of the main geological units before Gondwana amalgamation, the chapter focuses on the tectonic evolution of the Antarctic continent from its inclusion as part of the Gondwana supercontinent to the break-up of this landmass and the repositioning of Antarctica at southern polar latitudes since the Early Cretaceous. The geological evolution of the Antarctic continent is reviewed considering two main time periods: (i) c. 600–450 Ma, covering the processes which were active immediately before and during the amalgamation of Gondwana; and (ii) c. 450– 180 Ma, including all the major events that occurred after the ﬁnal stage of Gondwana amalgamation to the time immediately before the break-up phase. A subsequent section addresses the last 180 Ma during which present-day Antarctica and the other southern continents and surrounding oceanic basins formed as a consequence of the fragmentation of Gondwana. After a general overview of the most signiﬁcant plate tectonic stages and coeval magmatic products, the attention Corresponding author. Tel.: þ39-0577-233812; Fax: þ39-0577-233938;

E-mail: [email protected] (F.M. Talarico).

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is then devoted to one of the most investigated regions in Antarctica; the Transantarctic Mountains and the Ross Sea sector of the Western Antarctic Rift System. Some of the main persisting problems and potential research themes for on- and off-shore activities are discussed in the conclusions, which also include some of the most interesting palaeoclimatic issues that are essential to improving our understanding of the polar climate, ice ages and their inﬂuences on Earth’s climate system in the Cenozoic to present time.

7.1. Introduction The geological literature has for a long time demonstrated the fundamental role of Antarctica as a key continent for providing data and constraints to supercontinent reconstructions and the plate tectonic evolution of the Southern Hemisphere. Du Toit’s (1937) Gondwana reconstruction could be hypothesised on the basis of the results of the heroic period ﬁrst expeditions (i.e. the Glossopteris-bearing Beacon sandstones collected during the last Scott’s expedition). The Mesoproterozoic Rodinia supercontinent (Dalziel, 1991; Moores, 1991) provides another example with fundamental pieces of evidence linked to the central position envisaged for Antarctica with respect to the other continental blocks, albeit the conﬁgurations differ for the proposed supercontinent. Progressively revealed by increased research activities of national and multinational expeditions following the 1957–1958 International Geophysical Year to those planned for the International Polar Year 2007/2008, the geological record of Antarctica has now a similarly valuable interest and importance in the context of palaeoenvironmental and palaeoclimatic investigations, particularly those focused on the initiation of Cenozoic glaciation, the stability of the polar ice sheets and the complex interaction among tectonic, sedimentary and climatic processes. In the following sections, the Antarctic geological record is reviewed in the context of the major phases of the late Neoproterozoic–Cenozoic plate tectonic evolution of the Southern Hemisphere. The chapter describes the evolution of the Antarctic continent from its inclusion as part of the Gondwana supercontinent to the break-up of this landmass and the repositioning of Antarctica at southern polar latitudes since the Early Cretaceous (c. 120 Ma). The chapter also highlights some of the most interesting palaeoclimatic issues, which are considered essential to understanding the polar climate, ice ages and their inﬂuences on Earth’s climate system throughout the Cenozoic.

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The aim is to give a general overview which should be complementary to the more detailed information included in the accompanying chapters devoted to speciﬁc aspects of the geological record. Moreover, the review will highlight some of the main persisting open problems and those aspects of Antarctic tectonic evolution that are still not completely investigated or are under debate and may potentially be understood as major research themes for future on- and off-shore research.

7.2. The Present-Day Geotectonic Setting of Antarctica The Antarctic continent comprises three primary tectonic regions: (i) East Antarctica; (ii) West Antarctica and the associated West Antarctic Rift System (WARS); and (iii) the Transantarctic Mountains (Fig. 7.1). East Antarctica is thought to feature Precambrian continental lithosphere c. 35–45 km thick (Bentley, 1991), stable, coherent and topographically high (Cogley, 1984), that held a central position in the Palaeozoic supercontinent of Gondwana (Tingley, 1991) as it did in the Mesoproterozoic supercontinent Rodinia (Dalziel, 1991; Moores, 1991). In contrast, West Antarctica is an amalgamation of low-lying, 20–35 km thick, younger crustal blocks (Janowski and Drewry, 1981; Dalziel and Elliot, 1982). The West Antarctic Rift lithosphere compares to major Cenozoic continental rift systems (Behrendt, 1999). In the Ross Sea, the WARS borders the Transantarctic Mountains on their eastern side as a broad region of thinned continental crust associated with Cretaceous and episodic Cenozoic extension (Behrendt et al., 1991a,b; Behrendt, 1999). The WARS has high heat ﬂow (83–126 m W m2) (Blackman et al., 1987; Berg et al., 1989) and thick (5–14 km) sedimentary basins with recent faulting (Cande and Leslie, 1986; Cooper et al., 1987; Hamilton et al., 2001). The Ross Archipelago (LeMasurier and Thomson, 1990) is currently active with fumarolic activity associated with alkaline volcanism at Mt. Erebus and Mt. Melbourne. The crust in the WARS is currently B1972 km thick (Trehu, 1989; Cooper et al., 1997). The Transantarctic Mountains are approximately 2,500 km long and 200 km wide, dividing East Antarctica from West Antarctica with peaks that rise over 4 km above sea level. Crustal thickness estimates under the Transantarctic Mountains vary between 20 and 45 km (ten Brink et al., 1993, 1997; Cooper et al., 1997; Busetti et al., 1999; Kanao et al., 2002; Bannister et al., 2003). These mountains sharply differ from most mountain ranges of similar size and lateral extent because their formation did not reﬂect any compressional orogenic phases, but rather a process thought to have been

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Figure 7.1: Map showing the main geographic subdivision of Antarctica in three domains which also correspond to distinct tectonic regions: West Antarctica, Transantarctic Mountains and East Antarctica. The location of the two active volcanoes in the Ross Sea area is also shown (ME: Mt. Erebus; MM: Mt. Melbourne). Antarctic bed topography is based on the BEDMAP dataset (Lythe et al., 2000).

directly linked and causally related to the development of the WARS (e.g. ten Brink et al., 1997; Studinger et al., 2004, and references therein). The Transantarctic Mountains uplifted B6–10 km in an asymmetric tilt block formation and underwent denudation from the Cenozoic to the Cretaceous (Fitzgerald, 1992, 1995; Studinger et al., 2004). The later part of the uplift and denudation phases occurred under persisting spreading/extension within the western Ross Sea (Cande et al., 2000) and has been concomitant with voluminous sediment inﬁlling from the Late Eocene/Oligocene to present (Barrett et al., 2000, 2001; Hamilton et al., 2001).

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7.3. The Main Geological Units of Antarctica Before Gondwana Amalgamation As with all other continents, Antarctica consists of a number of Archean/ Early Proterozoic cratons (older than 1.5 Ga) surrounded by successively younger belts that formed and/or accreted to the continental margins, as products of convergent plate tectonic events such as subduction of oceanic crust underneath continental crust and/or collision of two former separated continents (Fig. 7.2). Cratons are conﬁned to East Antarctica and, at least two among several, are generally accepted: the small Grunehogna Craton (ca. 3.0 Ga basement covered by ﬂat-lying and undeformed sediments older than 1 Ga) interpreted as a fragment of the African Kalahari Craton, and the East Antarctic Craton s.s. The East Antarctic Craton s.s. is exposed only in limited areas including the southern Shackleton Range, the eastern Thiel Mountains (Ford, 1963), the Miller and Geologists Ranges of the central Transantarctic Mountains, near the continental margin in Enderby Land, in the southern Prince Charles Mountains, in the Vestfold Hills, in Wilkes Land, in Terre Ade´lie and in George V Land. The exposed basement, only occasionally covered by almost ﬂat-lying platform sediments (e.g. Thiel Mountains), consists of high-grade gneisses and granulite facies metamorphic complexes showing radiometric ages of older than 1,500 Ma up to just over 3,500 Ma in Enderby Land. Recently, parts of Enderby, Kemp and Mac. Robertson Lands have been separated from the East Antarctic Craton and attributed to younger orogenic belts, i.e. the Grenvillian and Pan-African belts s.l. (Fitzsimons, 2000a; Boger et al., 2001, 2002; Boger and Miller, 2004), so that northern Enderby Land is now considered a separate craton (Napier Complex). Moreover, the results of more recent investigations suggest that, in the next few years, our appreciation of the present conﬁguration of the East Antarctic Craton could change drastically and be subdivided in much smaller cratons including: (1) the Napier Craton (considered a fragment of Dwarhai Craton in India); (2) a craton comprising the southern Prince Charles Mountains, the southern Shackleton Range and perhaps the eastern Thiel Mountains; (3) the Vestfold Hills Craton; (4) the Mawson Craton (Fitzsimons, 2000a, 2003), probably consisting of the Miller and Geology ranges of the central Transantarctic Mountains, the shield areas of eastern Wilkes Land, Terre Ade´lie, George V Land, and their Australian counterpart, the Gawler Craton. Since the relationships between the Mawson Craton, the Vestfold Hills Craton, the southern Shackleton Range and the southern Prince Charles

Figure 7.2: Schematic geological map of Antarctica (modiﬁed from Kleinschmidt, 2007). The map shows the possible subice extent of the cratonic areas, the three ‘‘Grenvillian’’ belts, the Pan-African orogenic belts and the Ross Orogen. The map also shows the distribution of Beacon Supergroup and Ferrar Supergroup outcrops and of the two main Mesozoic orogenesis (Ellsworth or Weddell Orogeny and Antarctic Andean Orogen) which formed along the palaeo-paciﬁc margin of Gondwana in West Antarctica, and the three intra-plate fracture zones in East Antarctica (Lambert Graben or Lambert Rift, West Antarctic Rift System, Jutul Penck Graben and the Rennick Graben). Abbreviations – BLM: Bertrab, Littlewood, Moltke nunataks, Bu: Bunger Hills, Ele: Elephant Island, Ge: Geologists Range, JPG: Jutul Penck Graben, KS: Kuunga Suture, LHB: Lu¨tzow-Holmbukta, MF: Matusevich strike–slip fault, Mi: Miller Range, PCM: Prince Charles Mts., Pe: Pensacola Mts., RG: Rennick Graben, Sh: Shackleton Range, Sn: Snow Island, Sø: Sør Rondane, Th: Thiel Mts., Wi: Windmill Islands; Lands – DML: Dronning Maud Land, EL: Enderby Land, GVL: George V Land, KL: Kemp Land, KWL: Kaiser-Wilhelm-II-Land, MBL: Marie Byrd Land, MRL: Mac.Robertson Land, OL: Oates Land, PEL: Princess Elizabeth Land, TA: Terre Ade´lie, VL: Victoria Land, WL: Wilkes Land.

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Mountains are still totally unknown, a tectonic reconstruction is largely speculative and the cratonic areas could be even more reduced and further increased in number if, for example, 1.6–1.7 Ga-old orogens are separated. The orogenic belts separating the cratons as internal belts in East Antarctica or as elongated peripheral belts at the palaeo-Paciﬁc margin of Gondwana developed during ﬁve major orogenic cycles spanning in time from c. 1.3–0.9 Ga (Grenvillian-aged orogens), through 500–600 Ma (Ross and the Pan-African orogens), 250–200 Ma (Ellsworth or Weddell Orogen) and 150–90 Ma (Antarctic Andean Orogen) to the period between 50 Ma and recent, documented only by the active plate-margin processes in the area north-west of the Antarctic Peninsula. As in eastern North America and other continents, the Grenvillian-aged orogens in Antarctica are considered to record the fundamental orogenic event marking the amalgamation of the Rodinia supercontinent in Meso-/ Neoproterozoic times. Since polyphase deformation, high-grade metamorphism and magmatism resembles somewhat those in the cratons, the areas formed during the Grenvillian orogeny have been often undifferentiated from the cratons. Moreover, until the early 1990s, only one very long Antarctic Grenvillian orogen was assumed, following the Antarctic coast as a 250 km wide strip from Coats Land in the west up to George V Land in the east, occasionally specially named the ‘‘Circum East Antarctic Mobile Belt’’ (Yoshida, 1992). Such an extension has been demonstrated to be incorrect for Terre Ade´lie and George V Land and unproven for Coats Land, where the tiny outcrops represented by three little groups of nunataks (Littlewood, Bertrab and Moltke) consist of 1,100 Ma rhyolites and granophyres (Storey et al., 1994), but absolutely undeformed, thus not consistent with the existence of an orogenic belt. The existence of a Circum East Antarctic Mobile Belt is no more supported by new geological evidence which, instead, indicates the presence of three distinct Grenvillian-aged orogens (Fitzsimons, 2000a) (Fig. 7.2): (1) the Maud Belt, extended from western to central Dronning Maud Land to possibly parts of Sør Rondane, interpreted as the trace of welding the Grundhogna Craton and the Shackleton Range/southern Prince Charles Mountains craton; (2) the Rayner Belt (Yoshida and Kizaki, 1983) that in Enderby, Kemp and Mac. Robertson Lands separates the Napier craton from the southern Prince Charles Mountains craton, and was veriﬁed in the northern Prince Charles Mountains and the Rayner Complex (e.g. Boger et al., 2002); (3) the Wilkes province belt (Fitzsimons, 2000a) of Wilkes Land that is exposed in the Bunger Hills and Windmill Islands and merges the Vestfold Hills and the Mawson cratons.

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7.4. Antarctica in the Gondwana Supercontinent The ﬁt of the large continental blocks within Gondwana is well established (e.g. Lawver et al., 1998) as three continents, Australia, India and Africa, ﬁt well against the present-day rifted margins of Antarctica (Fig. 7.3). In contrast, due to considerable uncertainty over the number and position of some of the microplates within the West Antarctic region (e.g. Dalziel and Elliot, 1982; Storey, 1996; Storey et al., 1998), the Gondwana reconstruction is less well constrained along the Transantarctic margin of East Antarctica. In this region, the Ellsworth–Whitmore Mountains crustal block, Antarctic Peninsula, Thurston Island and Marie Byrd Land (subdivided into west and east portions, Divenere et al., 1996) are generally accepted as the main West Antarctic microplates, but other minor microplates (i.e. The Haag Nunatak, Berkner and Filchner microplates) are also used in reconstructions (e.g. De Wit et al., 1988). On the basis of palaeomagnetic data, geometry and geological constraints (particularly the presence of a large proportion of subduction-related rocks), the main West Antarctic microplates are generally retained in approximately their present location along the palaeo-Paciﬁc margin of Gondwana in most models (e.g. Lawver et al., 1992). To obtain their Gondwana pre-break-up positions, both the Antarctic Peninsula and the Thurston Island block (Grunow et al., 1991) are rotated anticlockwise from their present-day position with respect to East Antarctica, the two microplates of Marie Byrd Land are restored along a strike–slip fault (Divenere et al., 1995), and Campbell Plateau, Chatham Rise, and North and South Islands of New Zealand can be reconstructed to Marie Byrd Land.

Figure 7.3: (A) Gondwana tight ﬁt reconstruction (Early Jurassic). (B and C) Major initial stage in the break-up of Gondwana: (A) initial rifting stage (Late Jurassic); (B) change in the Gondwana break-up stress regime from dominantly north–south between east and west Gondwana to dominantly east–west with the two-plate system being replaced by a multiple-plate system (Early Cretaceous) (after Fitzgerald, 2002, with permission from the Royal Society of New Zealand). AP, Antarctic Peninsula; TI, Thurston Island; MBL, Marie Byrd Land; CR, Chatham Rise; CP, Campbell Plateau; SNZ, Southern New Zealand; NNZ, northern New Zealand; LHR, Lord Howe Rise; WS, Weddell Sea. Continent and microplate positions are from Lawver et al. (1992, 1998) with other information from Storey (1996).

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Detailed information about the different reconstructions and tectonic scenarios for Antarctica in Gondwana can be found in Lawver et al. (1998) and Larter et al. (2002). We will here focus on the major Antarctic geological constraints, key regions and datasets which represent a well-established basis for reconstructing the southern continents in Gondwana. In this context, our review will also face some controversial aspects, which are presently under debate, concerning the reconstruction of the main phases during the amalgamation of the supercontinent. The geological evolution of the Antarctic continent is reviewed considering two main time periods: (i) c. 600–450 Ma, covering the processes which were active immediately before and during the amalgamation of Gondwana; and (ii) c. 450–180 Ma, including all the major events that occurred after the ﬁnal stage of Gondwana amalgamation to the time immediately before the break-up phase. 7.4.1. Antarctica in the Late Precambrian–Early Palaeozoic (c. 600–450 Ma) Evolution of Gondwana The connections between Antarctica and the other southern hemisphere continents at the time of Gondwana are clearly documented by the so-called ‘‘Gondwanian’’ sequences which correlate with analogous stratigraphic successions in South America, Africa, India, Sri Lanka, Australia and New Zealand and by palaeontological evidence, such as the signiﬁcant ﬁrst appearance in Triassic time of the herbivorous reptile Lystrosaurus found in all Gondwanian continents, Antarctica included (Collinson, 1991; Retallack et al., 2005). Additional geological evidence for Gondwana’s reconstruction is provided by a number of older Antarctic geological provinces (such as the Archean Cratons and the Palaeozoic orogenic belts, e.g. the Ross– Delamerian Orogen) which ﬁt tightly across a closed Southern Ocean. In comparison to this well-established evidence, the reconstruction of tectonic models for the initial stages of the formation of Gondwana is more problematic. The amalgamation phase necessarily involved the aggregation of various continental fragments which derived from the fragmentation and dispersal of a former supercontinent – variously named Ur-Gondwana (Hartnady, 1991), Katania (Young, 1995), Palaeopangea (Piper, 2000) or Rodinia (McMenamin and McMenamin, 1990). However, the precise conﬁguration and modality of break-up of this supercontinent are yet not completely known and these uncertainties obviously propagate to the formulation of tectonic models for the constructive phase of Gondwana.

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Since the focus of this paper is on the Cretaceous–Cenozoic record, we will here avoid reviewing Rodinia models (e.g. SWEAT: Dalziel, 1991; Hoffman, 1991; Moores, 1991; AUSWUS: Karlstrom et al., 1999; see also reviews by Dalziel, 1997 and Meert and Torsvik, 2003) or alternative models (e.g. Palaeopangea, Piper, 1982, 2000) and we will brieﬂy summarize the main recent results in the reconstruction of the Late Precambrian–Early Palaeozoic tectonic evolution of Antarctica in Gondwana.

7.4.1.1. From Rodinia to Gondwana In spite of signiﬁcant uncertainties about the precise reconstruction of the global palaeogeography in Neoproterozoic time and the still incomplete geochronological framework, most authors agree that the break-up of Rodinia led to the development of extensive passive continental margins which are documented in the late Neoproterozic (c. 750–600 Ma) record of most present-day continents (Dalziel, 1991, 1992, 1997; Powell et al., 1993; Meert and Van der Voo, 1997; Cawood, 2005). In Antarctica, following Dalziel’s hypothesis (1991), the ﬁrst stage of Rodinia break-up involved a rifting phase which started at c. 750–725 Ma leading to the separation of Laurentian from the East AntarcticaþAustralia block and the formation of the intervening proto-Paciﬁc ocean. This process was accompanied by the drift of the cratonic blocks, presently exposed in South America and Africa (Amazonian, Rio de la Plata and Western African) (i.e. West Gondwana) which eventually collided with India, Sri Lanka and East Antarctica (East Gondwana). Most workers agree that key evidence of this evolution is stored in the Mozambique Belt (Holmes, 1951) or East African Orogen (Stern, 1994) and that this extensive orogenic belt formed as a result of the closure of a ‘‘Mozambique Ocean’’ and subsequent collision and amalgamation of East and West Gondwana during the Pan-African event (Dalziel, 1992; Stern, 1994; Shackleton, 1996). A review of structural and geochronological data from East Antarctica (Dronning Maud Land and Lu¨tzow Holm Bay) and comparison with the adjacent (in Gondwana) Falkland Microplate and south-eastern Africa led Jacobs and Thomas (2002) to corroborate the proposal by Jacobs et al. (1998) of a southward continuation of the Mozambique belt into Dronning Maud Land in Antarctica. A Mozambique suture zone in the Shackleton Range was suggested by Grunow et al. (1996) and conclusive evidence was found by Talarico et al. (1999) who described relics of ophiolites consisting of serpentinites and amphibolites with N-type MORB to OIB geochemistry and a maximum

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Sm–Nd age of c. 900 Ma. Metamorphic reworking of these rocks occurred under variable high P to medium P conditions in the eastern Shackleton Range; a stage of eclogite-facies metamorphism in the area has been recently reported by Schma¨dicke and Will (2006) in the central Shackleton Range. On the basis of these discoveries and thrust patterns in both the Shackleton Range and Western Dronning Maud Land, Kleinschmidt et al. (2002) proposed that the ophiolites may have formed part of a connection between the Palaeo-Paciﬁc and the Mozambique oceans, in the way the Drake Passage links the present Paciﬁc and Atlantic Oceans. The Mozambique Belt and its continuation in Antarctica shows a general N–S trending (Jacobs and Thomas, 2002). East of Dronning Maud Land, structural and petrological data indicate that the Lu¨tzow Holm Complex (Kriegsman, 1995) can be interpreted as representing the western end of a c. E–W, subperpendicular branch of the Mozambique Belt that might have extended to Prydz Bay, where Pan-African high-grade rocks have also been reported (Dirks and Wilson, 1995; Fitzsimons, 1997). The Lu¨tzow Holm–Prydz Bay Pan-African orogenic belt has provided sound evidence in contrast to the classical assumption that Eastern Gondwana formed during the consolidation of Rodinia in the Mesoproterozoic time and it remained tectonically stable until the modern continents rifted from Gondwana in the Mesozoic. The Lu¨tzow Holm–Prydz Bay–PanAfrican structures were initially considered part of a wider belt termed the Kuunga Orogen by Meert et al. (1995) or Kuunga Suture by Boger and Miller (2004), and interpreted as the result of the collision between East Antarctica (þAustralia) and India (þMadagascarþSri Lanka) at c. 535–520 Ma, after the amalgamation of India with the rest of Gondwana along the Mozambique suture (Fig. 7.4). More recent data actually suggest that there are three orogenic belts formed at about the same Pan-African period of 500–600 Ma in East Antarctica (Fig. 7.2): (1) the belt in the Shackleton Range–Dronning Maud Land–southern Sør Rondane–Lu¨tzow-Holmbukta region (‘‘East Antarctic Orogen’’ or ‘‘East Antarctic Belt’’ of Jacobs et al., 1998, renamed ‘‘Lu¨tzow Holm Belt’’ by Fitzsimons, 2000b); (2) the belt in the southern Prince Charles Mountains–Grove Mountains (‘‘Kuunga Suture’’ according to Boger et al. 2002); and (3) the belt in the Denman Glacier region, interpreted as prolongation of the Leeuwin Complex of Australia’s Pinjarra Orogen (e.g. Fitzsimons, 2000b). However, the exact extent of these orogens is doubtful because of the extensive ice cover, and this applies especially to the Denman Glacier belt. The Lu¨tzow Holm Belt is characterized by extensive thrust and nappe tectonics (Shackleton Range), by widespread and distinct late-orogenic collapse

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Figure 7.4: Main tectonic stages of the amalgamation of Gondwana (modiﬁed after Boger and Millar, 2004, with permission from Elsevier). Abbreviated cratons – D: Dwarhai craton; G: Gawler craton; K: Kalahary craton; P/Y: Pilbara/Yilgarn craton; sPCM: southern Prince Charles Mountains.

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structures (Shackleton Range), by thick molasse formations (Blaiklock Glacier Group: Shackleton Range; Urfjell Group: Dronning Maud Land), and by syn- and post-orogenic magmatism (Dronning Maud Land; Paech, 2004). Therefore, these Antarctic portions, formed during late Neoproterozoic to Cambrian, show typical characteristics of a collisional orogen. The three belts indicate the amalgamation of West Gondwana (South America, Africa and Grunehogna Craton) and East Gondwana (India, Australia and main East Antarctica). As proposed by Boger et al. (2001, 2002), Gondwana’s amalgamation may have taken place in two steps: the ﬁrst before 550 Ma and the second after 550 Ma. The ﬁrst step involved the amalgamation of West Gondwana with ‘‘Indo-Antarctica’’ (i.e. India and the northern Prince Charles MountainsþNapier Complex) documented by the Mozambique Belt and as its Antarctic prolongation, the Lu¨tzow Holm Belt. The second step led to the aggregation of these terranes to the rest of East Gondwana, i.e. the rest of East Antarctica and Australia and thus producing the Kuunga Suture. This model could explain the diacronous development of the Lu¨tzow Holm Belt, mainly older than 550 Ma in Dronning Maud Land and o550 Ma in the Shackleton Range, as well as the existence of an interleaved old alien element – the Grenvillian granophyres of Coats Land – maybe just an exotic terrane, or a mini-craton (Kleinschmidt, 2007).

7.4.1.2. The Antarctic record of the late precambrian–early palaeozoic evolution of the palaeo-paciﬁc margin of Gondwana The Transantarctic Mountains, at the margin of the East Antarctic shield, represent a key element in providing an important but still cryptic record of Proterozoic and Early Palaeozoic supercontinent history in the period from c. 800 to 500 Ma. The existing dataset and data that can be provided by future research are from sedimentary, plutonic and volcanic assemblages that potentially reﬂect different tectonic events integral to the Rodinian– Gondwanan transformation: from the break-up of Rodinia to the development of a transitional margin and plate-margin activity during Gondwanan assembly. In the Transantarctic Mountains, constraints on the timing of Rodinia break-up are provided by maﬁc magmatism and siliciclastic sedimentation of Neoproterozoic age from two areas. In the Skelton Glacier area, basalts interlayered with Skelton Group sediments yielded a Sm–Nd model age of 700–800 Ma (Rowell et al., 1993). In the Nimrod Glacier area (Cotton Plateau), gabbros and basalts interlayered with sediments of the Beardmore

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Group previously dated at 762 Ma (Sm–Nd isochron age, Borg et al., 1990) are now considered to have been emplaced at 668 Ma (U–Pb zircon age, Goodge et al., 2002). Thick siliciclastic successions were long interpreted as deep-water turbidities deposited in Proterozoic time along a rifted margin. Although the older parts of some successions may relate depositionally to the rifting process, recent investigations have demonstrated that some units are nearshore deposits and the depositional age of several major successions must be revised upward. For example, in the central Transantarctic Mountains, sandstones formerly included in the Beardmore Group and considered Neoproterozoic are now assigned to the Middle Cambrian or younger (Goodge et al., 2002). A transformation from drifting to active subducting mode is inferred for the late Neoproterozoic, starting at c. 700 Ma and continuing, most probably as a protracted contractional tectonic cycle including several discrete deformation events, with a tectonic climax during the Ross Orogeny, until the Ordovician (Goodge, 2001). The ﬁrst stage of the conversion from a passive margin to an active convergent margin is signalled by the structural inversion of cratonmargin sedimentary successions in central Transantarctic Mountains. Younger tectonic events between 560 and 480 Ma are generally referred to a broadly deﬁned Ross Orogeny, a long-lived tectonic process which developed along the active Gondwanan margin involving episodic deformation, calc-alkaline magmatism and syn-orogenic deposition of arcderived detritus in a sinistral-transpressive, continental-margin arc setting (Fig. 7.2). The orogenic belt is exposed from northern Victoria Land at the Paciﬁc end up to the Pensacola Mountains at the Atlantic end. Westernmost Marie Byrd Land (Edward VII Peninsula) has to be considered as part of the same orogenic belt, from which it became isolated only much later, during a major phase of the evolution of the WARS around the end of the Cretaceous. The Ross Orogen is characterized by folds, thrusts, very low- to high-grade metamorphism, granitoids, terranes and ﬂysch- and molasse-type sediments. Remarkable thrusts have been reported from Oates Land, which could be traced into the Australian continuation of the Ross Orogen, the Delamerian Orogen (Flo¨ttmann et al., 1993). The systematic distribution of high- and low-pressure types of metamorphism (e.g. Talarico et al., 2004) and of S- and I-type granitoids (e.g. Vetter and Tessensohn, 1987) led to the model of subduction of the palaeo-Paciﬁc beneath East Antarctica. The igneous activity occurred as early as 560–550 Ma ago and is considered to be related to initial subduction of palaeo-Paciﬁc lithosphere beneath the East Antarctic cratonal margin. In southern Victoria Land, the older plutons also include a peculiar suite of highly alkaline rocks (nepheline syenites and carbonatites)

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(Koettlitz Glacier Alkaline Province, Cooper et al., 1997) and a stage of Precambrian rift-related magmatism and sedimentation has been documented by Cook (2007) in the Skelton Glacier area. In northern Victoria Land, the Ross Orogen is made up of three so-called terranes: the high- to medium-grade and granite-dominated Wilson Terrane to the west, the low-grade turbiditic Robertson Bay Terrane to the east, and the low-grade and volcanic-rich Bowers Terrane in between. These terranes are considered to be allochthonous or just adjacent palaeogeographic domains including an intra-oceanic island arc (Bowers Terrane) and an accretionary wedge (Robertson Bay Terrane) (Tessensohn and Henjes-Kunst, 2005). In northern Victoria Land, a unique occurrence of ultra-high-pressure rocks, including well-preserved maﬁc eclogites as lenses and pods within metasedimentary gneisses and quartzites, decorates the tectonic boundary between the inboard Wilson Terrane and the Bower Terrane in the Lanterman Range. Geological, petrological and geochronological studies indicate that maﬁc, ultra-maﬁc and felsic host rocks in this region underwent a common metamorphic evolution with an eclogite facies stage about 500 Ma ago at temperatures of up to about 8501C and pressures greater than 2.6 GPa (Di Vincenzo et al., 1997; Palmeri et al., 2003, 2007). 7.4.2. Antarctica in the Upper Palaeozoic–Mesozoic (c. 450–180 Ma) Evolution of Gondwana In this time window, the main geological events so far known in the Antarctic geological record include the formation of a continental basin where a dominantly ﬂuvial succession, the Beacon Supergroup, was deposited in Devonian to Triassic times, and the development of two main deformational events generally known as the middle Palaeozoic Borchgrevink Orogeny and the Permo-Triassic Gondwanian Orogeny.

7.4.2.1. Borchgrevink Orogeny The occurrence in northern Victoria Land of a calc-alkaline magmatic suite of middle Palaeozoic (c. 350 Ma) age is considered as evidence of an orogenic event which is temporally well separated from the waning phase of the Ross Orogeny (Craddock, 1970). This suite includes dominant granitoids (Admiralty Intrusives) and minor felsic volcanic (Gallipoli Volcanics). The granitoids form several plutons which are mainly concentrated in the

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Robertson Bay Terrane, although some cut across the tectonic boundary between the Bowers and Wilson Terrane, where the coeval volcanics are also concentrated. All the intrusions are characterized by isotropic fabrics and discordant contacts with respect to the surrounding metasediments suggesting a post-tectonic emplacement. These features and meagre evidence of radiometric data for concomitant metamorphism and deformation have so far limited a comprehensive reconstruction of the tectonic setting and development of the Borchgrevink Orogeny in northern Victoria Land. Nevertheless, the continental arc geochemical signature of the Devonian– Carboniferous magmatic suite indicates that the plutonism most likely occurred as the result of a renewed period of subduction activity along the palaeo-Paciﬁc margin (Kleinschmidt and Tessensohn, 1987; Borg and DePaolo, 1991). Similar suites are known in Tasmania, New Zealand and the Campbell Plateau (Gibson and Ireland, 1996; Bradshaw et al., 1997) and in the Ford Range in Marie Byrd Land. A correlation between the Borchgrevink Orogeny of northern Victoria Land with the Tasmanian Orogeny was put forward by Findlay (1987). Elsewhere in Antarctica, marine sediments from the Ellsworth Mountains and Pensacola Mountains are considered to have been deposited in the same time window, within a basin that according to Elliot (1975) could have extended from the Weddell Sea to the Ross Sea.

7.4.2.2. Beacon Supergroup The Beacon Supergroup consists of dominantly continental sedimentary deposits which constitute a generally ﬂat-lying, 2.5–3.5 km thick cover developed over a marked unconformity (Kukri Peneplane) above the Ross orogenic belt throughout most part of the Transantarctic Mountains (Fig. 7.2). Similar sequences are known in limited outcrops in East Antarctica (Prince Charles Mountains, Dronning Maud Land and Ellsworth Mountains) (Barrett, 1991) and as bedrock of the Cenozoic glaciomarine sediments in the Victoria Land Basin (VLB) in the Ross Sea as documented by the CRP-3 drill-hole (Cape Roberts Science Team, 2000). Outside Antarctica, similar sedimentary rocks, collectively called Gondwanian sequences, occur in South Africa, Australia and South America. The deposition of these sediments in Antarctica started in Devonian time with the Taylor Group, consisting dominantly of quartz-arenites and conglomerates. Deposited as the result of erosion and ﬂuvial processes under arid and semiarid conditions (Campbell and Claridge, 1987), the Taylor Group accumulated in a series of basins along the palaeo-Paciﬁc margin of

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Gondwana. After the deposition of fossiliferous siltites, the subsequent stage corresponds to an erosional phase, probably related to a glacial event which led to the deposition of glaciogenic sediments in late Carboniferous–Early Permian time. In the Transantarctic Mountains, all the Carboniferous to Triassic sediments are formally known as the Victoria Group, which includes carbonaceous layers and feldspathic sandstones. The low sedimentation rate (c. 12.5 m/My) and absence of concomitant compressional deformation suggest that deposition occurred over a thick continental crust (Barrett, 1991) but different tectonic models have been proposed, including a passive margin (Isbell, 1999), intra-cratonic (Barrett, 1991; Woolfe and Barrett, 1995) or marginal/back-arc (Bradshaw and Webers, 1988).

7.4.2.3. Ellsworth or Weddell Orogeny The Ellsworth or Weddell Orogeny (or – in a larger context – also known as the Gondwanide Orogeny including components in southern Africa – Cape Fold Belt – and South America – Sierra de la Ventana Fold Belt) occurred in Permo-Triassic time ca. 250–200 Ma. As noted by Cawood (2005), this orogeny overlaps with the end Palaeozoic assembly of Pangea (Li and Powell, 2001), through ocean closure and accretion of Gondwana, Laurussia (LaurentiaþBaltica) and Siberia, as well as completion of terrane accretion in the Altaids. Stratigraphic and geochronological data (Dalziel, 1982; Dalziel and Elliot, 1982; Storey et al., 1987; Trouw and De Wit, 1999; Johnston, 2000) indicate that Permo-Triassic deformation of variable intensity and distribution occurs throughout West Antarctica and the adjoining Cape Fold Belt of southern Africa. In Antarctica, this orogenic event is well documented in the Ellsworth–Whitmore Mountains and in the Pensacola Mountains, where upright to inclined folds with axial planar cleavage are inferred to have formed in a dextral transpressive environment (Curtis, 1998) (Fig. 7.2). The Ellsworth–Pensacola Mountains chain represents the fold belt, while the Whitmore Mountains represent the magmatic arc of an Andean-type orogen. It merges with the Ross Orogen in the Pensacola Mountains, where it partly overprints the older Ross-aged structures. Elsewhere, deformation is heterogeneously distributed, with Storey et al. (1987) noting that in the Antarctic Peninsula, unconformities previously ascribed to the Gondwanide Orogeny are younger and that the only event related to Gondwanide deformation s.s. is regional metamorphism at 245 Ma of parts of the Trinity Peninsula Group.

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The Ellsworth Orogen trends noticeably crossways to the palaeo-Paciﬁc margin of Antarctica as indicated by the Ross Orogen. This obliqueness is due to secondary rotation, as proven by palaeomagnetic investigations by Funaki et al. (1991) and conﬁrmed by Randall and MacNiocaill (2004).

7.4.2.4. The Antarctic Andean Orogen The orogen of the Antarctic Andes occupies the entire Antarctic Peninsula down to the Walgreen Coast (Fig. 7.2). It formed mainly in three episodes: (i) Late Jurassic through Early Cretaceous (150–140 Ma), (ii) mid-Cretaceous (B105 Ma) and (iii) BTertiary (B50 Ma to recent), and is partly still active (e.g. Birkenmajer, 1994; Vaughan and Storey, 1997). Thus, it represents the youngest growth zone of the continent. The Antarctic Andes are a typical subduction orogen accompanied by orogenic magmatism in the form of granitic plutons and volcanic rocks. In detail, the deformation and metamorphism are very complicated, because they are polyphase. Folding and thrust faulting is reported mainly from the southern portion of the Antarctic Peninsula (Palmer Land and Alexander Island) and from the extreme north (Trinity Peninsula and eastern South Shetland Islands). The distribution of related metamorphism is also heterogeneous, including high-pressure metamorphism with blueschists characteristic of subduction complexes, e.g. on Elephant Island (Trouw et al., 1991).

7.4.2.5. The plate-tectonically active parts of Antarctica The only plate-tectonically active part of Antarctica is situated north-west of the Antarctic Peninsula, in the South Shetland Islands (from Snow Island in the south-west up to Elephant Island in the north-east) and the Bransﬁeld Strait (Fig. 7.2). North-west of the South Shetland Islands, a small section of the ocean ﬂoor, called the Drake Plate (i.e. the remnant of the older, but largely subducted Phoenix Plate), is being subducted at the South Shetland Trench beneath the Antarctic Plate. Related, mainly andesitic, volcanism forms the island arc of the South Shetland Islands. Parts of the South Shetland Islands (part of Livingston Island and Elephant Island) belong – as does the Peninsula itself – to earlier stages of the Antarctic Andean orogeny and consist of strongly deformed Jurassic trench sediments.

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The Bransﬁeld Strait is located south of the subduction-related volcanic island arc and forms an active extensional basin accompanied by tholeiitic volcanism, partly submarine, partly as active island volcanoes (e.g. Penguin Island and Deception Island; Smellie and Lo´pez-Martı´ nez, 2002). The Bransﬁeld Strait often is regarded as a classic example of a back-arc basin, but recently this has been disputed (Gonza´les-Casado et al., 2000).

7.5. Antarctic Record of Gondwana Break-Up and Dispersal of the Southern Hemisphere Continents This section addresses the 180 Ma to recent time window during which present-day Antarctica and the other southern continents and surrounding oceanic basins formed as a consequence of the fragmentation of Gondwana, and the tectonic processes involved in the drift and dispersion of the various continental fragments. After a general overview of the most signiﬁcant plate tectonic stages and coeval magmatic products, the attention is devoted to one of the most investigated regions in Antarctica, the Transantarctic Mountains and the Ross Sea sector of the Western Antarctic Rift System.

7.5.1. Summary of Main Tectonic Stages of Gondwana Break-up The ﬁrst major tectonic stage in the break-up of Gondwana corresponds to an initial rifting phase that started in the Weddell Sea, initially as a back-arc basin, in the Late Jurassic (Fig. 7.4) (Lawver et al., 1991). This stage involved right-lateral transtension as East Gondwana (Antarctica, Australia, India and New Zealand) and West Gondwana (South America and Africa) moved apart with stretching beginning in the north and propagating southward (Lawver et al., 1992). Initial break-up involved complex geodynamic evolution characterized by the rotation and translation of several microplates, such as the Ellsworth Mountains block, a displaced part of the Gondwanide fold belt. The original position of the various microplates is still controversial, but the Ellsworth–Whitmore Mountains crustal block most likely originated from near south-eastern Africa (e.g. Curtis and Storey, 1996). Rotation of West Antarctic microplates must have been accomplished by c. 165 Ma (the time of opening of the Weddell Sea), and rotation of the Ellsworth–Whitmore Mountains crustal block was ﬁnished before translation into its present position by 175 Ma (Grunow et al., 1987). Rotation of microplates did not apparently involve the production of oceanic crust

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(Marshall, 1994) but may have occurred as block rotation with controlling faults concealed beneath Mesozoic sedimentary basins in the Weddell Sea (Storey, 1996). The break-up has been explained by several authors as the result of a hot mega-plume, which, according to Storey and Kyle (1999), could have promoted domal uplift and formation of a triple junction. According to Dalziel et al. (1999), the Gondwana plume may have also caused or expedited formation of the early Mesozoic Gondwanide fold belt, due to the buoyancy of a hot plume acting on the downgoing slab and causing it to ﬂatten. In both cases, plume activity led to the production of magma batches reﬂecting different degrees of plume–lithosphere interaction, which migrated along crustal shear zones to ultimately form the various large igneous provinces. The plume-related magmatic products are represented by huge withinplate maﬁc and felsic magmatic provinces in many Gondwana continents as well as Antarctica (Cox, 1988; White and MacKenzie, 1989; Storey, 1996). In the Transantarctic Mountains (LeMasurier and Thomson, 1990), Jurassic maﬁc rocks are generally known as the Ferrar Supergroup or the Jurassic Ferrar Large Igneous Province (FLIP). They are divided into a volcanic component, the Kirkpatrick Basalt – preceded by extensive phreatomagmatic volcanoclastic rocks (Elliot et al., 2006; Viereck-Go¨tte et al., 2007) – and the intrusive dolerite (diabase) sills and dikes of the Ferrar Dolerite. Geochemically, the FLIP is unusual with upper crustal-like characteristics such as high large ion lithophile element concentrations and enriched isotopic signatures (87Sr/86SrW0.709). The crust-like signature suggests derivation from an enriched lithosphere, possibly connected to subduction along the Paciﬁc margin of Gondwana during the Palaeozoic and Mesozoic. The maﬁc rocks are concentrated within a long linear belt exposed in Tasmania, Antarctica and South Africa. Cox (1988) considered that the linear pattern of the Ferrar and Tasman provinces could not be compatible with classic circular plumes and proposed a hot line rather than a hot spot. A number of rifts that intersected at the Dufek Massif (Elliot, 1992) could have favoured the development of zones of weakened lithosphere, which acted as pathways for lateral migration of magmas derived from lithospheric sources. In spite of the still not completely understood tectonic setting, it is important to note that, similarly to other continental ﬂood basalt provinces, all the maﬁc products formed during a short period of eruption (Tasman province: 175718 Ma; Ferrar Supergroup: 180–183 Ma; Dronning Maud Land province: 17772 Ma; Karoo province: 18272 Ma) (Hergt et al., 1989; Hooper et al., 1993; Heimann et al., 1994). Coeval felsic intrusions considered to be rift-related are known in West Antarctica (Storey et al., 1988) and southern South America (Chon-Aike or Tobı´ fera province) (Gust et al., 1985).

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The second major geodynamic stage occurred in the Early Cretaceous (Fig. 7.4), as a consequence of the change of the Gondwana break-up stress regime from dominantly north–south between East and West Gondwana to dominantly east–west with the two-plate system being replaced by a multiple-plate system (e.g. Lawver et al., 1992). This modiﬁcation in stress regime is thought to have induced large-scale ductile deformation concentrated along shear zones in the Antarctic Peninsula (Storey et al., 1996) and thin-skinned deformation and inversion of existing sedimentary basins such as the Latady Basin (Kellogg and Rowley, 1989). By c. 110 Ma, the microplates of West Antarctica had nearly reached their present location with respect to East Antarctica. In the same time period, separation also began between India and Antarctica (Lawver et al., 1991). Initial stretching between Australia and Antarctica began as early as 125 Ma (Stagg and Willcox, 1992), but sea-ﬂoor spreading was delayed to c. 95 Ma (Cande and Mutter, 1982; Veevers et al., 1990; Royer and Rollet, 1997), in the Ross embayment (Elliot, 1992), as well as extension between the Lord Howe Rise and northern New Zealand (Lawver et al., 1992). By the Late Cretaceous, Antarctica had reached its ﬁnal polar location and conﬁguration, and the ﬁnal stage of break-up was completed when New Zealand (Campbell Plateau) rifted from Marie Byrd Land at 84 Ma (e.g. Stock and Molnar, 1987; Lawver et al., 1991). In this geodynamic context, four particularly large and conspicuous intraplate fracture zones have been investigated and all related to extensive and prolonged extensional regimes spanning in time from Mesozoic to present (Fig. 7.2). These major extensional zones include: the Lambert Graben or Lambert Rift (East Antarctica); the WARS and its main part, the Ross Sea Rift (Paciﬁc sector); the graben of Jutulstraumen and Penckmulde (occasionally called Jutul Penck Graben; Atlantic sector); the Rennick Graben as the main element of a strike–slip fault system in Victoria and Oates Lands. The Lambert Graben developed in the East Antarctic Craton and is ﬁlled by sediments of the Permo-Triassic Beacon Supergroup. Faulting started already during the early Palaeozoic, reached its peak in the Permian and continued to the Early Cretaceous (Hofmann, 1996). Possibly, subglacial Lake Vostok belongs to the same rift system, but somewhat offset. The continuation of the Lambert Graben is the Indian Mahanadi Rift south-west of Calcutta in the state of Orissa (Hofmann, 1996), ﬁlled with sediments of the same type and age as the Lambert Graben. The reconstruction of the Gondwanan India–Antarctica ﬁt by these graben systems coincides with

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reconstructions by Archean to Early Proterozoic elements for Rodinia. That means, interestingly enough, that the relationships of Antarctica and India during Rodinian and Gondwanan times do not differ substantially. The Ross Sea Rift is extremely wide (about 1,000 km). Its subsidence started during the late Mesozoic (about 140 Ma ago), reached its main activity in the Early Tertiary (about 40 Ma ago) and produced an enormous relief at its western shoulder. The difference in altitude between the tops of the Transantarctic Mountains and the ﬂoor of the adjacent Ross Sea exceeds 14 km. The crustal extension is combined with alkaline intra-continental volcanism, which is still active at Mt. Erebus (3,794 m) and at Mt. Melbourne (2,732 m), both located in Victoria Land (Kyle and Cole, 1974; Wo¨rner et al., 1989; Kyle, 1990a,b; Tessensohn and Wo¨rner, 1991). The Jutul Penck Graben of western Dronning Maud Land originated probably around 140 Ma ago or a little bit later (Jacobs and Lisker, 1999). The graben marks the boundary of the Grunehogna Craton towards the south-east and thus it follows a much older geological structure. A possible continuation of the Jutul Penck Graben into the still active East African rift system is under discussion (Grantham and Hunter, 1991). By no means is this out of the question, because parts of the East African rift system were already active in the Jurassic (Ring and Betzler, 1993). The strike–slip fault system of Victoria and Oates Lands runs obliquely to the Ross Sea Rift and is cut by it. The principle element of the system is the Rennick Graben, which is presently active, as demonstrated by earthquakes in 1952, 1974 and 1998. The graben contains downfaulted Ferrar volcanics and sediments of the Beacon Supergroup, which have been spectacularly folded and squeezed onto the graben shoulders (Rossetti et al., 2003). This demonstrates alternating dextral transpression and transtension (with formation of pull-apart basins), being parts of a complicated strike–slip system in detail (Rossetti et al., 2003). The 1974 quake happened some 120 km to the west at the parallel structure of the Matusevich Glacier. There is an ongoing discussion, as to whether this strike–slip fault system constitutes the continuation of oceanic fracture zones between Australia and Antarctica (e.g. the Tasman Fracture Zone) into the continental crust of Antarctica (Salvini et al., 1997; Salvini and Storti, 2003; Kleinschmidt and La¨ufer, 2006). 7.5.2. Tectonic Evolution and Magmatism in the Ross Sea Sector during the Cenozoic With a similar form to the Weddell Sea, the Ross embayment is one of the most striking morphological expressions of the WARS, a region of thin, and,

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by inference, extended continental crust whose regional boundaries are difﬁcult to deﬁne precisely and may have been different for different episodes of extension (Fig. 7.5). Three major episodes of extension have been proposed, with rifting starting in the Middle Jurassic, in coincidence with the onset of Gondwana break-up and the associated Middle Jurassic magmatism (Ferrar Group), and with subsequent episodes in the Cretaceous and late Cenozoic. The active rifting may continue to the present as suggested by the active volcanoes along the western margin of the Ross Sea. However, most extension and thinning of the Ross Sea crust is considered to have taken place during the Mesozoic rift period and the tectonic connection between the Ross and Weddell embayments is largely unclear (Behrendt et al., 1992; LeMasurier, 2007, and references therein). In particular, our knowledge of the Jurassic episode is problematic since, although the location of rifting is well constrained in the South Africa – Queen Maud Land region, neither the location nor the amount of extension is well known in the Transantarctic Mountains and Ross Sea. By comparison, the Cretaceous episode is better documented (LeMasurier, 2007, and references therein) and is considered to be closely related to the initial stage of stretching and rifting in the Weddell Sea region, where disruption of the plate margin involved the rotation and translation of the several crustal blocks forming the present West Antarctica, including the Ellsworth/Whitmore Block, which was translated to the boundary between the Ross and Weddell embayments. Since movement of the West Antarctic crustal blocks was largely completed by 110 Ma, later tectonic activity in the WARS has been restricted to the Ross sector of the TAM and the margin of the Ellsworth–Whitmore Mountains, with the continuation of the TAM into the Weddell Sea region considered a remnant Jurassic rift (Schmidt and Rowley, 1986; LeMasurier, 2007). The ﬁrst direct evidence of rifting in the Ross embayment is documented in Marie Byrd Land where a voluminous suite of maﬁc dikes and anorogenic A-type granites (the so-called ‘‘Byrd Coast Granites’’) have been dated as early as 10775 to 102–95 Ma (Storey et al., 1999). Considering Bradshaw’s (1989) suggestion that a sea-ﬂoor spreading centre was subducted beneath Mesozoic New Zealand around 105 Ma, the production of the maﬁc dykes could have been linked to the subduction of such a spreading centre at 10775 Ma, followed by the emplacement of the anorogenic silicic rocks at 102–95 Ma. On the basis of the age of Anomaly 33 (83 Ma) identiﬁed off Campbell Plateau, the earliest sea-ﬂoor spreading between Campbell Plateau and Marie Byrd Land must be B85 Ma. The unequivocal reconstruction of

Figure 7.5: (A) Map of the Transantarctic Mountains and West Antarctic Rift System (after Fitzgerald, 2002, with permission from the Royal Society of New Zealand). Structural features along the TAM and in the West Antarctic Rift System are from Tessensohn and Wo¨rner (1991), Fitzgerald (1992), Fitzgerald and Baldwin (1997) and Salvini et al. (1997). DAZ: Discovery accommodation zone from Wilson (1999). Basin positions in the Ross Sea (light grey) are from Davey and Brancolini (1995) and include North Basin (NB), Eastern Basin (EB), Central Trough (CT) and the Terror Rift (TR) lying within the Victoria Land Basin (VLB). The approximate outline of the Wilkes subglacial basin is shown inboard of the TAM in a dot pattern. West Antarctic microplates are after Dalziel and Elliot (1982); see Fig. 7.3 for abbreviations. (B) Generalized crustal proﬁle A–A (in A) across the Ross Sea is after Cooper et al. (1991). Early rift grabens lie beneath regional unconformity U6 (heavy line). Late rift-faults and intrusive rocks deform the VLB and some small basement grabens. Note intra-basement reﬂector (dashed line) on eastern edge of VLB, possibly a detachment fault extending under the entire basin. DSDP site 270 is projected onto this crosssection.

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the Campbell Plateau against Marie Byrd Land and the remarkable match of the deduced ocean–continent boundary as well as the alignment of the eastern edge of the Eastern Ross Basin with the northern edge of the Campbell Basin on the Campbell Plateau are evidence that the entire Eastern Ross Basin should have experienced extension prior to initiation of the Paciﬁc–Antarctic Ridge at B85 Ma. The Paciﬁc–Antarctic spreading centre played a major role in the region in late Cenozoic time. With its initiation, a triple junction developed with three extensional arms (Australian–Antarctic spreading centre, Australian–Paciﬁc arm and the Paciﬁc–Antarctic spreading). Geochronological constraints are provided by early sea-ﬂoor spreading anomalies identiﬁed between Tasmania and the South Tasman Rise. Key tectonic events include (i) the jumping of the Australia–Antarctica spreading centre about Anomaly 30 (B65 Ma) to between the South Tasman Rise and northern Victoria Land, (ii) the completion of sea-ﬂoor spreading in the Tasman Sea at Anomaly 24 time (B54 Ma) as a consequence of the reformation of the ridge–ridge–ridge triple junction as a new spreading centre that went between Campbell Plateau and the Tasman Sea anomalies and (iii) the opening of the B40 km wide Adare Rift within ‘‘older’’ sea-ﬂoor off the western edge of the Ross Sea sometime after Anomaly 13 (B33 Ma) (Cande et al., 2000; Cande and Stock, 2006). Satellite gravity data suggest that the Adare Rift steps east, and then is transformed westward along the southern margin of the Coulman High to the region of the Terror Rift, a comparable structure in width, located along the western edge of the VLB in the western Ross Sea. Striking products of the large-scale Cenozoic geodynamic evolution in the Ross Sea area are four major sedimentary basins which developed by rifting of previously thinned continental crust, possibly along pre-existing crustal faults, and coincide with major localized crustal thinning (to less than 10 km), and, in the case of the VLB (the westernmost one), with high heat ﬂow and alkaline volcanism (Blackman et al., 1987; White, 1989; Cooper et al., 1991; Behrendt et al., 1993; LeMasurier, 2007). Extensive late Cenozoic volcanism is also inferred from aeromagnetic data. These basins include from E to W: the Eastern Basin – underlying most of eastern Ross Sea; the Central Trough – running north–south discontinuously through central Ross Sea; the Northern Basin – underlying the north-eastern Ross Sea margin; and the VLB underlying south-western Ross Sea, adjacent to the Transantarctic Mountains. The region has been extensively investigated through a number of geophysical surveys. But, due to the lack of appropriate age data for the main seismostratigraphical units, inferences on the timing of Cenozoic rifting in the entire Ross Sea remain largely speculative. Based on the results of the

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Cape Roberts Drilling Project (Barrett et al., 2001), the VLB is now considered to be mostly late Eocene or younger in age. Recent drilling on the western margin of this basin has indeed indicated an onset of subsidence at about 34 Ma, signiﬁcantly younger than the onset of uplift of the adjacent Transantarctic Mountains (55 Ma) (Fitzgerald, 1992) and thus identifying an issue in the relationship of the uplift of rift margin mountains to the subsidence of the adjacent rift basin. However, the effect of uplift episodes at this margin on the sedimentary sequences recorded may need to be considered. The fact that the basin extends only from the Ross Island region to Terra Nova Bay, whereas the Transantarctic Mountains continue further north and south, would indicate that the basin may have originated from other processes in addition to simple extension. A transtension or ‘‘pull-apart’’ process has been suggested, limited by transfer mechanisms associated with the igneous activity recorded in the north by the ‘‘Polar Three’’ anomaly off Coulman Island and in the south by the magnetic anomalies south of the Ross Sea Fault (Bosum et al., 1989). Another signiﬁcant geological component in the Ross Sea area tectonic evolution is represented by an extensive alkalic magmatic province, one of the largest in the world and including two active volcanoes (Mt. Erebus at Ross Island and Mt. Melbourne on the Ross Sea coast in northern Victoria Land; LeMasurier and Thomson, 1990). These rocks, either exposed or suggested by aeromagnetic studies in the area under the Ross Sea and West Antarctic Ice Sheet (Behrendt et al., 2002; Bell et al., 2006), occur on either side of the WARS in Marie Byrd Land (West Antarctica) and in the western Ross Sea. In the western Ross Sea, the oldest rocks are restricted to coastal areas in northern Victoria Land where they consist of Eocene to Oligocene alkali intrusive rocks (Meander Intrusive), interpreted as the eroded remnants of subvolcanic magmatic complexes (Mu¨ller et al., 1991; Rocchi et al., 1999, 2002). Elsewhere, the alkaline magmatism is predominantly volcanic and has been subdivided into the informal, but geographic and petrologic distinct, Hallett, Melbourne and Erebus volcanic provinces (Kyle, 1990b). Two small isolated occurrences of basalts dated at 16–20 Ma occur around the head of the Scott Glacier (Stump et al., 1980). In the Erebus volcanic province (Kyle, 1990a), there is a continuous eruptive sequence from 19 Ma to the present day, with the Mt. Erebus crater showing a persistent anorthoclase phonolite lava lake. The oldest exposed rocks of this province occur on the northern slopes of Mt. Morning where 19 Ma trachyandesites lavas are intruded by 16–18 Ma trachyte dikes. Volcanic ash (tephra) layers in CRP drill cores demonstrate evidence for older Miocene trachytic eruptions. The tectonic setting suggests that the southern extent of the volcanism may be controlled along a major transfer

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fault, which coincides with the southern boundary of the Terror Rift (Wilson, 1999). Many of the larger volcanic centres have a radial distribution around Mt. Discovery or Mt. Erebus. The radial distribution is interpreted as a result of upwelling of a mantle plume (the Erebus plume), which was centred under Mt. Discovery prior to 4 Ma and then migrated to its present position under Mt. Erebus (Kyle, 1990a). 7.5.3. Tectonic Evolution of the Transantarctic Mountains The Transantarctic Mountains, one of the major young mountain chains on Earth, separate East Antarctica from the West Antarctica Rift System and the grounded East Antarctic Ice Sheet from the marine-based West Antarctic Ice Sheet over a substantial portion of the Antarctic interior along the margin of the Ross embayment (Fig. 7.5). The regional structural architecture of the Transantarctic Mountains remains poorly known in most regions because of the extensive ice cover. The East Antarctic Ice Sheet hides the structure of the mountains along the Polar Plateau, preventing the identiﬁcation of the extent of mountain structures into East Antarctica. In the McMurdo Sound coastal area of the Ross embayment, thin but extensive piedmont glaciers obscure the structural boundary (‘‘Transantarctic Mountains Front’’) with the off-shore Victoria Land rift basin. The main drainage for the East Antarctic Ice Sheet into the Ross embayment of West Antarctica is through outlet glaciers carved through the mountains. It has long been inferred that these outlet glaciers developed where faults cut transverse to the mountain trend (e.g. Gould, 1935; Gunn and Warren, 1962; Grindley and Laird, 1969; Davey, 1981; Wrenn and Webb, 1982; Cooper et al., 1991; Tessensohn and Wo¨rner, 1991; Fitzgerald, 1992), and the occurrence of pseudotachylites has been proved a valuable tool to constrain the age of some faults (Di Vincenzo et al., 2004). In most cases, however, there is little direct evidence either for the existence of a fault or of its age. This is evidently a key issue to be addressed in order to determine if appropriate structures are present to allow differential uplift of discrete mountain blocks (e.g. van der Wateren et al., 1999), to understand the localization of valley incision and how erosion has contributed to mountain uplift, and to provide constraints on the development of the pathways for drainage of the East Antarctic Ice Sheet. In contrast to other young mountain belts, such as the Alpine–Himalayan system and the North American and Andean Cordilleras, which formed at convergent plate boundaries, there is a tight genetic link between the uplift of the Transantarctic Mountains and intra-plate processes associated with the

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rifting within the Antarctic plate. These mountain ranges are generally interpreted as a high-relief rift ﬂank uplift (van der Wateren et al., 1999), which occurred during the Mesozoic–Cenozoic break-up of the Gondwana supercontinent (Cooper et al., 1987, 1991; Tessensohn and Wo¨rner, 1991; Davey and Brancolini, 1995; Fitzgerald and Stump, 1997). The development of the Transantarctic Mountains in an extensional, rather than a contractional, tectonic regime was already recognized by pioneering Antarctic geologists, who interpreted the mountain chain as a fault-bounded horst block (David and Priestley, 1914; Gould, 1935). More recent structural investigations indicate that the Transantarctic Mountains consist of a linear to curvilinear chain of asymmetric tilt blocks bounded on the West Antarctic edge by a major normal fault zone and subdivided by transverse faults (Fitzgerald et al., 1986; Tessensohn and Wo¨rner, 1991; Fitzgerald, 1992; Tessensohn, 1994a,b; Fitzgerald and Baldwin, 1997). Active rift tectonism and mountain uplift have been inferred from the presence of active volcanism and Neogene–Quaternary age faulting in the western portion of the rift and the Transantarctic Mountains (Behrendt and Cooper, 1991; Davey and Brancolini, 1995; Jones, 1997). The Cenozoic–Cretaceous asymmetric uplift and subsequent erosion exposed basement rock and older sediments along the coastward side of the Transantarctic Mountains, leaving younger sediments only on the inland side. Apatite ﬁssion track analysis in the McMurdo Sound area indicate an uplift of B6 km since c. 55 Ma, though other sectors of the Transantarctic Mountains record denudation events in the Late Cretaceous also (Fitzgerald, 1992, 1995; Studinger et al., 2004) (Fig. 7.6) and in the Early Cretaceous (e.g. Scott Glacier area: Stump and Fitzgerald, 1992; Fitzgerald and Stump, 1997). The cause of this uplift and denudation is the subject of continuing debate, the reconstruction of Cenozoic tectonic processes being complicated by the complex interplay between a number of factors including the regional plate geodynamics, rifting style, erosion rates, subsidence and formation of thick sedimentary layers, the volcanic activity and the glacial processes. The possible mechanisms for the uplift include thermal buoyancy due to conductive or advective heating from the extended upper mantle of the hotter West Antarctic lithosphere (Stern and ten Brink, 1989; ten Brink and Stern, 1992), simple shear extension (Fitzgerald et al., 1986), isostatic rebound due to stretching of the lithosphere through normal faulting (Bott and Stern, 1992), plastic necking (Chery et al., 1992), elastic necking (van der Beek et al., 1994) and rebound response to erosion (Stern and ten Brink, 1989). In the McMurdo Sound area, some of these mechanisms are based on speciﬁc assumptions about the crustal and upper mantle structure beneath the ‘‘Transantarctic Mountains Front’’, as well as about the timing of the

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Figure 7.6: Schematic diagram showing the variation of exhumation events along the TAM at different localities (after Fitzgerald, 2002, with permission from the Royal Society of New Zealand). SCG, Scott Glacier area; BDM, Beardmore Glacier Area; SHG, Shackleton Glacier area; SVL, Southern Victoria Land; TNB, Terra Nova Bay; NVL, Northern Victoria Land. A relative scale only is shown for exhumation as the amount of exhumation at any one locality will vary across the range. Early or late Cretaceous exhumation is not always present throughout an area (e.g. Scott Glacier region; see Fitzgerald and Stump, 1997).

rift-related processes in the nearby VLB. Some constraints have already been provided by gravity studies (e.g. Davey and Cooper, 1987; Reitmayr, 1997), and by seismic reﬂection data (O’Connell and Stepp, 1993; Della Vedova et al., 1997). Data from the ACRUP seismic experiment indicate thickening crust beneath the Transantarctic Mountains to a depth of 38 km, and quite low P-wave velocities (7.6–7.7 km/s) in the mantle beneath the VLB (Della Vedova et al., 1997), while slow S-wave velocities are inferred at 60–160 km depth from surface wave analysis (Bannister et al., 1999), suggesting that the upper mantle is anomalously warm at that depth. More recently, the ﬁndings of drilling projects in the McMurdo Sound area (CIROS, Cape Roberts Project) have resolutely constrained the age of

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the onset of subsidence at the westernmost margin of the VLB. The ﬁrst direct geological evidence of a major pre-Oligocene uplift phase of the Transantarctic Mountains comes from the oldest strata cored in the CIROS-1 and CRP-3 drill-holes (Barrett et al., 1989, 2001). These include granitic clasts eroded from exposed basement to the west, implying that the Transantarctic Mountains were at least half of their present height by then, for erosion had cut through the more than 2,000 m of Devonian–Jurassic Gondwana cover beds to basement (Barrett et al., 1989, 2001). In the Cape Roberts drill core, the presence of the Devonian Arena Formation (Beacon Supergroup) as bedrock beneath the Cenozoic sediments indicates that signiﬁcant uplift and unrooﬁng of the Transantarctic Mountains must have occurred prior to the Oligocene (Barrett et al., 2001).

7.6. Open Problems and Potential Research Themes for Future Geoscience Investigations in Antarctica In contrast to all other continents, 99% of Antarctica (including its rocks and geological structures) is covered by a major ice sheet. Less than 1% of the continent provides our geological knowledge and even this has not been investigated thoroughly in all places. On the other hand, about half of the covered area has been surveyed geophysically, through mainly aeromagnetic and gravimetric techniques. Therefore, further investigations are needed and the image of the Antarctic geological structure and its history will likely change, improve and be completed in the future. Several mountain ranges in the Antarctic are difﬁcult to access and therefore have been seldom visited in the past. However, their study would decidedly improve the understanding of geologic and geotectonic connections. It is evident that there are still conspicuous gaps in our knowledge of Marie Byrd Land, the Pensacola Mountains, eastern Dronning Maud Land and East Antarctica between 601 and 1201E. Moreover, the unknown ice-covered interior needs targeted studies starting from well-known areas, for example, tracing known geological rock complexes from their exposed areas under the ice with the help of suitable geophysical methods. The conﬁdence we can have in geophysical interpretations declines with increasing distance from directly accessible rock complexes. Even a few spot checks of rock samples from isolated deep drill-holes would provide a better reliability of interpretations, i.e. calibration of airborne geophysical data. The concept of the formation of supercontinents represents an important step in higher level global cycles in order to unravel the history of

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development of continents and oceans, and to improve our knowledge on global plate tectonic and geodynamic processes. Since Antarctica, and particularly East Antarctica, had a central position in both Rodinia, between 1,300 and 700 million years ago, and Gondwana, between 550 and 200 million years ago (today’s southern continents), geological information archived in this ‘‘keystone’’ region is therefore important not only for a wellfounded analysis of local conditions, but also in relation to our understanding of Earth system processes in general. Abundant evidence exists for the former connection of both Antarctic cratonic areas (the East Antarctic and Grunehogna cratons) to geologically similar provinces on neighbouring continents. The study of East Antarctica is a prerequisite for the reconstruction of the assembly and break-up processes of both supercontinents, Rodinia and Gondwana. This pertains to each of the plate conﬁgurations, including the orogens as well as the processes in general, which have led to the assembly of the continental lithosphere. Geological and palaeomagnetic data show that East Antarctica includes at least two older cratonic fragments, and Grenvillian and Pan-African structures in coastal exposures support this notion. The larger part (East Antarctic Craton s.s.) continues into India and Australia, while a small fragment (Grunehogna Craton) connects to the Kalahari Craton of Africa. Newest results show that even the East Antarctic Craton s.s. consists of a number of cratonic nuclei (Fitzsimons, 2000a; Boger and Miller, 2004). These results conﬁrm the importance of Antarctica for the reconstruction of continental distribution in early Earth history. Signiﬁcant contributions to the understanding of global plate tectonic and geodynamic processes are also stored in the orogenic belts of Pan-African age (600–500 million years ago) which contain important information about the juxtaposition of West and East Gondwana. In Antarctica, the Shackleton Range and parts of East Antarctica between Dronning Maud Land, Lu¨tzow Holm Bukta and Prydz Bay belong to these belts (Buggisch et al., 1990; Jacobs et al., 1998; Tessensohn et al., 1999; Fitzsimons, 2000b; Boger et al., 2002; Paech, 2005; Buggisch and Kleinschmidt, 2007). Equivalent rocks have been found in the African Mozambique Belt (Paech, 1985; Jacobs et al., 1998; Paech et al., 2005). So far, nothing is known about how these fragments continue under the ice and how they can be connected with the better researched mountains in Dronning Maud Land and the Transantarctic Mountains. Equally unknown is the formation, age and relation of subglacial mountains in the East Antarctic interior, e.g. the Gamburtsev Subglacial Mountains. At approximately the same time as the development of internal Pan-African orogenic belts, the geographic domain corresponding to the present Transantarctic Mountains experienced the Ross orogeny, the

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result of dominant accretionary tectonic processes produced by the subduction of the palaeo-Paciﬁc plate under the palaeo-Paciﬁc Gondwana margin. There is widespread consensus in considering the end of the Ross orogeny as a nearly synchronous event along the different sectors of the belt (e.g. Stump, 1995; Encarnacion and Grunow, 1996). Voluminous granitoids (Granite Harbour Intrusive Complex) intruded as batholiths at c. 530– 480 Ma (Encarnacion and Grunow, 1996) represent a unifying feature throughout the length of the Transantarctic Mountains (Borg et al., 1990; Stump, 1995). But although the general tectonic history of the Ross Orogen is fairly well known within each of the major segments of the Transantarctic Mountains, signiﬁcant variations in lithostratigraphic, structural and metamorphic patterns, as well as in granitoid geochemical afﬁnity, are evident between the different segments. Considerable uncertainty still remains about the onset of the subduction, the tectonic setting of the early granitoids with variable chemical afﬁnities (from calc-alkaline to alkaline and carbonatite) of southern Victoria Land, the nature of the contact between the orogenic belt and the East Antarctic Craton and the relations with the Pan-African structures of the Shackleton Range. Until our knowledge of the relationships of the tectono-metamorphic histories and of the detailed chronology of the magmatic episodes between the segments is improved a comprehensive tectonic model of the development of the Ross improved remains to be formulated. The importance of Antarctica is also well recognized with regard to the destructive processes of plate tectonics, including the fragmentation of supercontinents. The young continental margins and rift structures of Antarctica, as well as their development, document the break-up of Gondwana and its meaning for the present-day habitats and environmental conditions. Aside from a small section along the Antarctic Peninsula, the continental margins of the Antarctic depict the fault structures of the breakup of Gondwana leading to the formation of the present southern continents and oceans. Antarctica has lain in its south polar position since at least the Late Cretaceous (approximately 130 Ma). Its isolation from the neighbouring continents through the opening of the Southern Ocean and the formation of the Antarctic Circumpolar Current began at this time. The further break-up process led to the formation of the Southern Ocean. The exact opening processes and their effects on palaeoceanography have not yet been satisfactorily reconstructed because high-quality magnetic data are lacking in many key areas of the oceans. This is especially the case for the South Paciﬁc and areas between Antarctica and Africa and Australia, respectively.

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For an understanding of the plate tectonic development of the Antarctic continent, as well as climatic and biological developments, several events must be considered: 130–100 million years ago: the opening of the Weddell, Lazarev and Riiser-Larsen Seas and thus the opening of the southern Atlantic and Indian Oceans; 110–80 million years ago: the genesis of the southern Kerguelen Plateau; 80–40 million years ago: the separation of Tasmania/Australia from Antarctica; approximately 40 million years ago: the main phase of development of the Ross Sea Rift; 30–20 million years ago: the separation of South America from Antarctica, and thus opening of the Drake Passage and Scotia Sea with the formation of the Circumpolar Current. Aside from the basic questions about the mechanisms and the consequences of these global processes, e.g. of climatic nature, the breakup of Gondwana also led to the current mosaic of continental plates. It was the starting point for many current processes and patterns, for example, the glaciation of the polar regions, the present-day oceanic and atmospheric circulation, the distribution of climatic zones and biota, and ﬁnally the global conditions for our existence. The present state of knowledge indicates that the distinct steps of the break-up of Gondwana must be determined more precisely. In particular, the phase since the total isolation of Antarctica is globally relevant for the development of ocean circulation and climate patterns and it must therefore be further deciphered. More speciﬁcally, an increased research effort would be essential to provide new and more complete on- and off-shore (mainly from drill-holes) data to ﬁll the numerous knowledge gaps on the Cenozoic Antarctic glacial history and to allow a better timing of the Antarctic Circumpolar Current onset (Barker and Thomas, 2004; Barker et al., 2007). Indeed, the important role of the polar regions is emphasized by general circulation models (Sloan et al., 1996; Oglesby, 1999). The presence of thick ice sheets in the continents inﬂuences latitudinal gradients and sea-ice formation, and it drives the formation of cold Antarctic bottom water which through deep ocean currents reaches the lower latitudes. The oxygen isotope record from deep-sea cores, and eustatic changes inferred from sequence stratigraphic records on passive continental margins, have led our knowledge of the long-term and broad history of the Antarctic Ice Sheet. However, interpretations based on these proxy records of glacio-eustasy have little

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direct conﬁrmation from geologic records in Antarctica, and in numerous cases have led to conﬂicting interpretations (Harwood et al., 1991, 1993; Moriwaki et al., 1992; Wilson, 1995; Miller and Mabin, 1998). In this context, since direct data from the Antarctic region are essential to validate these models, a series of drilling projects have targeted the Antarctic continental margin to retrieve high-resolution stratigraphic records of Antarctica’s glacial and climatic history (Cooper and Webb, 1992; Barker et al., 1998; Hambrey et al., 1998; Barrett et al., 2000, 2001; Hambrey, 2002) and more work in this area is planned for the next decade e.g. (ANDRILL, Harwood et al., 2002). Palaeoclimatic reconstructions based on the results of the Cape Roberts drilling project in the Ross Sea region for the early Oligocene to early Miocene time show two major climatic phases both documenting a warmer climate in the Antarctic than that which prevails today (Hambrey et al., 1998; Barrett et al., 2000, 2001; Naish et al., 2001). Also in the early Quaternary, which was characterized especially in the northern hemisphere by cool climatic conditions, the Cape Roberts core proved that there were times when the Antarctic experienced higher temperatures than today (Hambrey et al., 1998). However, despite the successes of drilling to date, there remain major unresolved questions concerning Cenozoic tectonic and palaeoenvironmental evolution of the Antarctic region. Chief among these is the still controversial problem of how stable the Antarctic Ice Sheets were during the last 20 My and the timing of the onset of glaciation. The former has been one of the main scientiﬁc targets of the ﬁrst ANDRILL drilling seasons in 2006 and 2007 austral summers (Harwood et al., 2002) and the ongoing investigations are attempting to determine the responses of past ice sheets and shelfs to climate forcing, including variability at a range of time-scales over the last 20 My. The latter has remained an elusive target for both CIROS and Cape Roberts drilling. Indeed, in CRP-3, Oligocene strata passed via an unconformity into the pre-glacial Devonian–Triassic Beacon Supergroup rocks. The search goes on for sites which will recover an Eocene and earlier Cenozoic, or even a Cretaceous record. This issue, along with others, will likely be addressed by future drilling programs that will also complement data from the Ross Sea with data from other coastal regions of Antarctica, and will provide a deeper insight into the climatic relevance of this region.

ACKNOWLEDGEMENTS This chapter is a contribution to the SCAR programme Antarctic Climate Evolution (ACE). Acknowledgements for funding and logistical and

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technical support with the ﬁeld campaigns and drilling projects in the McMurdo Sound area are made to the US Ofﬁce of Polar Programs (NSFOPP), Antarctica New Zealand, the Italian National Antarctic Program (PNRA), the Alfred Wegener Institute for Polar and Marine Research (AWI) and the Deutsche Forschungsgemeinschaft (DFG). Valuable comments by F. Florindo and M. Siegert, and careful corrections and reviews by C.A. Ricci and E. Stump are also much appreciated.

REFERENCES Bannister, S., Snieder, R. K., & Passier, M. L. (1999). Shear-wave velocities under the Transantarctic Mountains and terror rift from surface wave inversion. Geophys. Res. Lett., 27, 281–284. Bannister, S., Yu, J., Leitner, B., & Kennett, B. L. N. (2003). Variations in crustal structure across the transition from west to east Antarctica, southern Victoria Land. Geophys. J. Int., 155, 870–884. Barker, P. F., Barrett, P. J., Camerlenghi, A., Cooper, A. K., Davey, F. J., Domack, E. W., Escutia, C., Kristoffersen, Y., & O’Brien, P. E. (1998). Ice sheet history from Antarctic continental margin sediments: The ANTOSTRAT approach. Terra Antartica, 5(4), 737–760. Barker, P. F., Filippelli, G. M., Florindo, F., Martin, E. E., & Scher, H. D. (2007). Onset and role of the Antarctic circumpolar current. Deep-Sea Res. II, 54, 2388–2398. Barker, P. F., & Thomas, E. (2004). Origin, signature and palaeoclimatic inﬂuence of the Antarctic circumpolar current. Earth Sci. Rev., 66, 143–162. Barrett, P. J. (1991). The Devonian to Triassic Beacon Supergroup of the Transantarctic Mountains and correlatives in other parts of Antarctica. In: R. J. Tingey (Ed.). The Geology of Antarctica. Oxford Monographs on Geology and Geophysics. Clarendon Press, Oxford, pp. 120–152. Barrett, P. J., Davey, F. J., Ehrmann, W. U., Hambrey, M. J., Jarrard, R., van der Meer, J. J. M., Raine, J., Roberts, A. P., Talarico, F., & Watkins, D. K. (Eds). (2000). Studies from the Cape Roberts Project, Ross Sea, Antarctica, Scientiﬁc Results of CRP-2/2A, Parts I and II. Terra Antartica, 7 (4/5), 665 pp. Barrett, P. J., Hambrey, M. J., Harwood, D. M., Pyne, A. R., & Webb, P. N. (1989). Synthesis. In: P. J. Barrett (Ed.). Antarctic Cenozoic History from the CIROS-1 Drillhole, McMurdo Sound. DSIR Bull., 245, pp. 241–252. Barrett, P. J., Ricci, C. A., Bucker, C. J., Davey, F. J., Ehrmann, W. U., Laird, M. G., van der Meer, J. J. M., Raine, J. I., Smellie, J. L., Talarico, F., Thomson, M. R. A., Verosub, K. L., & Villa, G. (Eds). (2001). Studies from the Cape Roberts Project, Ross Sea, Antarctica, Scientiﬁc Report of CRP-3, Parts I and II. Terra Antartica, 8 (3/4), 620 pp.

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Developments in Earth & Environmental Sciences, 8 F. Florindo and M. Siegert (Editors) r 2009 Elsevier B.V. All rights reserved DOI 10.1016/S1571-9197(08)00008-6

Chapter 8

From Greenhouse to Icehouse – The Eocene/Oligocene in Antarctica J. E. Francis1, S. Marenssi2, R. Levy3, M. Hambrey4,, V. C. Thorn1, B. Mohr5, H. Brinkhuis6, J. Warnaar6, J. Zachos7, S. Bohaty7 and R. DeConto8 1

School of Earth & Environment, University of Leeds, Leeds, LS2 9JT, UK Instituto Anta´rtico Argentino, Universidad Buenos Aires and CONICET, Buenos Aires 1010, Argentina 3 Geosciences, University of Nebraska Lincoln, Lincoln, NE 68588, USA 4 Institute of Geography and Earth Sciences, Aberystwyth University, Ceredigion SY23 3DB, Wales, UK 5 Humboldt-Universita¨t zu Berlin, Museum fu¨r Naturkunde, D-10099 Berlin, Germany 6 Palaeoecology, Department of Biology, Faculty of Sciences, Laboratory of Palaeobotany and Palynology, Utrecht University, Budapestlaan 4, 3584CD Utrecht, The Netherlands 7 Earth and Planetary Sciences Department, University of California, Santa Cruz, CA 95064, USA 8 Department of Geosciences, University of Massachusetts, Amherst, MA 01003, USA 2

ABSTRACT The change from a warm, ice-free greenhouse world to the glacial Antarctic icehouse occurred during the latest Eocene–earliest Oligocene. Prior to this, during the Early–Middle Eocene, Antarctica experienced warm climates, at least on the margins of the continent where geological evidence is present. Climates appear to have been warm and wet, the seas were warm and plants ﬂourished in a frost-free environment, although there is some suggestion of valley glaciers on King George Island. Climate signals in the geological record show that the climate then cooled but not enough to allow the existence of signiﬁcant ice until the latest Eocene. Glacial deposits on Seymour Island indicate that ice was present there at Eocene/Oligocene boundary times. Further south, ice-rafted clasts in drill cores Corresponding author. Tel.: +44(0)1970 621860;

E-mail: [email protected] (M. Hambrey).

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from the Ross Sea region and deposits of tidewater glacier origin in Prydz Bay conﬁrm the presence of ice at the continental shelf by the earliest Oligocene. This matches the major Oi-1 oxygen isotope event in the marine record. On land, vegetation was able to persist but the warmth-loving plants of the Eocene were replaced by shrubby vegetation with the southern beech Nothofagus, mosses and ferns, which survived in tundra-like conditions. Throughout the Oligocene, glaciation waxed and waned until a major glacial phase in the Miocene. Coupled climate–ice sheet modelling indicates that changing levels of atmospheric CO2 controlled Antarctica’s climate. Factors such as mountain uplift, vegetation changes and orbital forcing all played a part in cooling the polar climate, but only when CO2 levels reached critical thresholds was Antarctica tipped into its icy glacial world.

8.1. Introduction One of the most intriguing challenges in Antarctic Earth history is to understand the fundamental climate change from the past greenhouse world with no major polar ice caps to our present icehouse that is dominated by the vast ice sheets on the Antarctic continent. This change across a major climate threshold holds many clues that will help us understand the potential changes our world may undergo in future. Geological evidence from rocks and fossils from the Antarctic continent and from marine oxygen isotopes that record changes in temperature, ice volume and water masses indicate that ice sheets built up on Antarctica from about Eocene/Oligocene (E/O) boundary times, approximately 34 million years ago. This is named the Oi-1 event in the marine oxygen isotope record and is represented by the appearance of several glacial deposits in the rock record. However, the actual pattern of climate cooling and the causes of glaciation are far from understood, and there are hints of the presence of ice in the Late Eocene. Indeed, it is possible that ice existed on Antarctica even during the Cretaceous (Miller et al., 2005; Tripati et al., 2005). This chapter reviews our current understanding of the greenhouse– icehouse transition in Antarctica. It covers the interval of climate change from the warm greenhouse climates of the Early Eocene through to the ﬁrst appearance of ice and the establishment of glacial conditions during the Oligocene. Several lines of evidence are presented for climate and environmental change: the sedimentary rock record on the continent provides clues to the nature of the cooling climate during the Eocene; latest

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Eocene and Oligocene sediments recovered from marginal basins in drill cores contain the earliest undisputed glacial deposits; fossil plants and palynomorphs from both the continent and marginal basins have yielded information about cooling climates in the terrestrial realm; marine microfossils hold clues to ocean circulation and the signiﬁcance of ocean gateways; the marine isotope record of the open oceans tells us about changes in deep ocean temperatures as the result of climate cooling and ice growth; and ﬁnally, the cause of cooling and the birth of the icehouse world is explored through computer modelling. The ﬁnal section summarizes our current understanding of the greenhouse–icehouse transition in Antarctica.

8.2. Climate Signals from the Sedimentary Record Sedimentary strata of Eocene and Oligocene age are exposed in the Antarctic Peninsula region and provide an important record of environmental conditions on land and in shallow marine settings during this period. In addition, there are intriguing hints about environments at higher latitudes, extracted from erratic boulders composed of Eocene and Oligocene fossiliferous sediments in the Ross Sea region, derived from sub-glacial outcrops. A summary of the important outcrops and their environmental signal is given below. Sedimentary rocks of Palaeogene age are exposed around the northern part of the Antarctic Peninsula, on the South Shetland Islands and on Seymour Island (Fig. 8.1). The sediments were deposited in very different tectonic settings and environments – the South Shetland sequence is of terrestrial volcanic and sedimentary deposits that represent an outer-arc (Birkenmajer, 1995) or fore-arc (Elliot, 1988) succession; the sequence on Seymour Island consists of marine clastics deposited in a back-arc basin, the uppermost beds of a regressive megasequence (Hathway, 2000). Both contain evidence of Palaeogene cooling and the ﬁrst appearance of ice but in different settings.

8.2.1. Antarctic Peninsula Region The Palaeogene back-arc deposits exposed on Seymour Island and Cockburn Island comprise more than 1,000 m of shallow marine to coastal fossiliferous clastic sedimentary rocks mainly of Palaeocene and Eocene age (Elliot, 1988; Sadler, 1988; Marenssi et al., 1998). The Maastrichtian–Danian sequence forms a simple BN–S homocline dipping gently to the east while the Late Palaeocene and the Eocene sediments ﬁll incised valleys trending NW–SE.

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Figure 8.1: Map of the Antarctic continent (A) and the Antarctic Peninsula (B) showing some of the locations mentioned in the text. The Early Eocene to latest Late Eocene La Meseta Formation (Elliot and Trautman, 1982) is an unconformity-bounded unit (La Meseta Alloformation of Marenssi et al., 1996, 1998) of maximum composite thickness of 720 m, which ﬁlls a 7 km wide valley cut down into older strata after the regional uplift and tilting of the Palaeocene and Cretaceous older beds. The La Meseta Formation comprises mostly poorly consolidated siliciclastic ﬁne-grained sediments deposited in deltaic, estuarine and shallow marine environments as part of a tectonically controlled incised valley system (Marenssi, 1995; Marenssi et al., 2002). The richly fossiliferous Eocene sediments have yielded the only fossils of land mammals in the whole Antarctic continent (Reguero et al., 2002) along with fossil wood, fossil leaves, a rare ﬂower, plus marine vertebrates (including giant penguins) and invertebrates (Stilwell and Zinsmeister, 1992; Gandolfo et al., 1998a,b,c; Francis et al., 2004; Tambussi et al., 2006). Overall, evidence from fossils, sediments and geochemistry from Seymour Island indicates generally warm and ice-free conditions during the earliest part of the Eocene but followed by gradual cooling. Dingle et al. (1998), based on chemical analysis (chemical index of alteration, CIA) and clay mineralogy, recorded a climatic deterioration from warm, non-seasonally wet conditions during the early Middle Eocene (smectite-dominated clay assemblage and CIA values o0.7 to W0.6) to a latest Eocene cold, frostprone and relatively dry regime (illite-dominated clay association and CIA

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values o0.6). Gazdzicki et al. (1992) showed a 6 m decrease in d13C values in biogenic carbonates and Tatur et al. (2006) recorded an increase in the Cd/Ca ratio in bivalve shells from the upper part of the La Meseta Formation. They interpreted these results as a change from stratiﬁed to vigorously mixed oceanic conditions related to the cryosphere development in the Southern Ocean by the end of the Eocene. Unpublished stable isotope (d18O and d13C) measurements (Feldmann and Marenssi) obtained from molluscan shells of the La Meseta Formation show a period of warmer seawater temperatures during 51–47 Ma and a drop of 1.51C during 35–34 Ma. Further, Ivany et al. (2008), also undertaking isotopic analysis of bivalves from the La Meseta Formation, propose the possibility of winter sea-ice formation in this region during the late Middle Eocene (c. 37 Ma). They suggest mean ocean temperatures have approached freezing at this time by assuming a sea water d18O value of 3m as predicted for the east Antarctic Peninsula by Huber et al. (2003) in their early Eocene climate model. This information is consistent with ice-rafted debris reported for the upper part of the La Meseta Formation (Doktor et al., 1988) and palaeoclimatic evidence of a severe climatic deterioration towards the end of the Eocene. It is possible that by the end of the Eocene, limited ice, perhaps as valley glaciers, was already present in the area. By the end of the Eocene, it is possible that an ice sheet extended over much of the peninsula, although average Seymour Island Shelf temperatures did not reach below zero (Ivany et al., 2008). Ivany et al. (2006) reported 5–6 m thick glacial deposits that conformably overlie the typical marine sandstones of the La Meseta Formation but are beneath the glacial diamictites of the younger Weddell Sea Formation. Based on dinocyst stratigraphy and strontium isotopes, these authors suggested an age of 33.57–34.78 Ma for these glaciomarine deposits (supported by marine Sr dates from Dingle and Lavelle, 1998; Dingle et al., 1998; Dutton et al., 2002), at or very close to the E/O boundary. There is therefore a small but intriguing window into the early stages of the icehouse world in the James Ross Basin. 8.2.2. King George (25 de Mayo) Island, South Shetland Islands The South Shetland Islands contain a different story of early glacial events within a fore-arc/outer-arc terrestrial setting. King George Island and neighbouring Nelson Island consist of several tectonic blocks bounded by two systems of strike–slip faults of Tertiary (54–21 Ma) age (Birkenmajer, 1989). Thus, considerable differences in

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stratigraphic succession, age and character of the rocks occur between particular blocks. The stratigraphic sequence includes mainly Late Cretaceous to Early Miocene island-arc extrusive and intrusive rocks comprising mainly terrestrial lavas, pyroclastic and volcaniclastic sediments often with terrestrial plant fossils. Hypabyssal dykes and plutons intrude the latter. Fossiliferous marine and glaciomarine sediments are also represented that provide clues to palaeoclimates. Several sequences of tillites crop out within these complicated sequences, representing glacial and interglacial events. Reports of supposed Eocene-age tillites at Magda Nunatak (Birkenmajer, 1980a,b), named the Krakow Glaciation and dated at 49 Ma, have been disproved by Sr dating (Dingle and Lavelle, 1998). However, Birkenmajer et al. (2005) describe valley-type tillites between lava sequences in the E/O Point Thomas Formation in Admiralty Bay. K–Ar dating gives ages of 41–45 Ma for the lavas below the tillites and 45–28 Ma in the lavas above. Birkenmajer et al. thus propose that a glacial period occurred at 45–41 Ma during the Middle Eocene, being the oldest record of alpine glaciers in West Antarctica. A clear record of glacial activity is present as diamictites and ice-rafted deposits within the Polonez Cove Formation, of mid-Oligocene age (26–30 Ma) (Troedson and Smellie, 2002). This is called the Krakowiak Glacial Member. At its maximum extent, ice was grounded on a shallow marine shelf. Interestingly, exotic clasts within this sequence may represent ice-rafted debris that was derived from as far away as the Transantarctic Mountains, suggesting marine-based glaciation in the Weddell Sea region. Non-glacial sediments that overlie the Polonez Cove Formation signal an interglacial period in the Late Oligocene (26–24.5 Ma), before another glacial phase in the Miocene. The only other terrestrial evidence for Oligocene ice, possibly representing local alpine glaciation, is from Mount Petras in Marie Byrd Land, West Antarctica, where deposits indicate volcanic eruptions beneath ice (Wilch and McIntosh, 2000).

8.2.3. Eocene Environments in the Ross Sea Region There is an intriguing record of Eocene pre-glacial environments in the Ross Sea region, around McMurdo Sound. Although the location of the Eocene outcrop is presumably under the ice sheet, there are no major exposures. A glimpse of Eocene environments is, however, provided by several hundred erratic boulders and cobbles recovered from coastal moraines around the shores of Mount Discovery, Brown Peninsula and Minna Bluff (Fig. 8.1).

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These fossiliferous glacial erratics (called the McMurdo Erratics), recovered from moraine around the northwest coast of Mount Discovery and Minna Bluff in southern McMurdo Sound, provide a window into the environment that may have existed along the coast of the gradually rising Transantarctic Mountains during the Eocene. The erratics are most likely derived from sub-glacial basins, such as Discovery Deep, that lie along the coast of the Transantarctic Mountains or basement highs situated to the east of the discovery accommodation zone (Wilson, 1999; Wilson et al., 2006). The erratics were distributed into their distinctive pattern of terminal and lateral retreat moraines during relatively recent advance and retreat of grounded ice into southern McMurdo Sound (Wilson, 2000). Subsequent basal adfreezing and surface ablation has transported the erratics to the surface of the McMurdo Ice Shelf. Although currently out of their original stratigraphic position, this suite of erratics provides us with a mechanism to obtain geologic data that are otherwise buried beneath the Antarctic Ice Sheets and fringing ice shelves. The McMurdo Erratics comprise a range of lithotypes and ages. Eocene rocks contain a rich suite of fossil ﬂoras and faunas including marine and terrestrial palynomorphs, diatoms, ebridians, marine vertebrates and invertebrates, terrestrial plant remains and a bird humerus. Biostratigraphic data from dinoﬂagellate cyst (dinocyst), ebridian and mollusc assemblages recovered from many of the erratics indicate that the majority of fossiliferous rocks range from Middle to Late Eocene, B43–34 Ma. Erratics collected between 1993 and 1996 (Stilwell and Feldmann, 2000) include several hundred samples of Oligocene, Miocene and Pliocene sediment. Although relatively rich dinocyst assemblages have been described from Oligocene Miocene Sequences from the Cape Roberts Cores, assemblages in post Eocene erratics comprise few taxa (typically o5 species). This general paucity of dinocyst species in late Palaeogene and Neogene sequences is observed in several other sites from the southern high latitudes (Wilson, 1989; Mao and Mohr, 1995) and may be a reﬂection of geographical and thermal isolation of Antarctica (McMinn, 1995; Williams et al., 2004) or preservation. The majority of the Eocene erratics record a suite of lithofacies that were deposited in coastal–terrestrial to inner shelf marine environments (Levy and Harwood, 2000a,b). These sediments were probably deposited within fan deltas that formed along the rugged coastline of the rapidly rising Transantarctic Mountains. Abundant macroinvertebrate faunas, including bivalves, gastropods, scaphopods, cirripeds, bryzoans, decapods and brachiopods, indicate that many of these sediments were deposited in a spectrum of predominantly shallow marine environments. The presence of terrestrial plant material and palynomorphs also suggests that the

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majority of the rocks were formed in nearshore environments. However, the occurrence of open marine dinocyst species and the absence of benthic diatom taxa in many of the ﬁne-grained lithofacies indicate that outer shelf open marine environments were also present in the source region. The Eocene erratics contain no direct or unequivocal sedimentological evidence for the presence of ice close to the basins in which the sediments were originally deposited. It is notable that erratics composed of diamictites recovered from the coastal moraines are all Oligocene and younger in age. Although rare, fossil leaves, wood and pollen recovered from several erratics provide a glimpse of the Eocene climate for the region. One erratic contains wood and leaves from Araucaria and Nothofagus trees, which suggests that the climate was not extreme. Cool temperate conditions with some winter snow may have been possible but temperatures were probably not cold enough to allow extensive ice at sea level (Francis, 2000; Pole et al., 2000). Spore and pollen assemblages recovered from the erratics reﬂect Nothofagus-podocarpaceous conifer-Proteaceae vegetation with other angiosperms growing in temperate climate conditions (Askin, 2000). Oligocene and younger erratics show a major drop in species richness, which is also noted in sequences recovered in CIROS-1 and the Cape Roberts Project (CRP) cores (Mildenhall, 1989; Raine and Askin, 2001). Fossil invertebrate remains recovered from the erratics include a humerus shaft from a pseudodontorn (giant bony-toothed sea bird) (Jones, 2000), a probable crocodile tooth (Willis and Stilwell, 2000) and teeth from two species of shark (Long and Stilwell, 2000). The small but signiﬁcant record of East Antarctic invertebrate fauna indicates a temperate to cool temperate marine environment. 8.2.4. Climate Evidence from Drilling on the Antarctic Margin The onset of glaciation in Antarctica is not yet well constrained, largely because no cores have yet been obtained that unequivocally provide a continuous transition from no-ice to ice-sheet scale glaciation. This is because (i) most cores terminated before the base of the glacigenic sediments was reached, (ii) a hiatus exists at the base of the glacigenic strata, as in CRP-3 in the Ross Sea (Barrett and Ricci, 2001a,b) and Ocean Drilling Program (ODP) Site 1166 in Prydz Bay (Shipboard Scientiﬁc Party, 2001a,b,c,d), or (iii) the age models based on multiple criteria have been revised several times (e.g. CIROS-1; Wilson et al., 1998a). On land, there is evidence for ice proximal-fjordal sedimentation possibly dating back to Oligocene time in the Prince Charles Mountains (Hambrey and McKelvey,

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2000a; McKelvey et al., 2001) and some exposures of the Sirius Group in the Transantarctic Mountains could also be this old (Sugden and Denton, 2004).

8.2.4.1. Drill cores in the western Ross Sea – CIROS-1 and Cape Roberts Two drill cores have been recovered from the western Ross Sea that approach or even cross the E/O boundary (Fig. 8.2). Drilling was undertaken from a sea-ice platform in spring-time, and was characterized

Figure 8.2: Location of the drill sites in the western Ross Sea. CIROS-1 and CRP-3 are those that bear on the E/O question (from Hambrey et al., 2002). Reproduced with permission of The Geological Society Publishing House, Bath, UK.

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by exceptionally high recovery (up to 98%). In the 702 m deep CIROS-1 hole (Barrett, 1989; Barrett et al., 1991), the lower part of the core was originally regarded as Late Eocene, with a breccia passing up into mudstone and sandstone. The boundary with the Oligocene was originally placed at about 570 m (Barrett et al., 1989), but magnetobiostratigraphic data (Wilson et al., 1998a) suggest that the E/O boundary is much higher at about 410–420 m. In either case, there is no obvious lithological transition, these ﬁner grained facies including alternations of weakly stratiﬁed sand and mud, with intraformational conglomerate and occasional diamictite. These facies are strongly bioturbated and contain exotic clasts, as well as intraclasts, while some beds are graded. Moving up core, a major hiatus exists at 366 m, which coincides with the Early/Late Oligocene boundary. Above is a suite of ﬁning-upwards sandstone beds, followed by alternating massive and stratiﬁed diamictites and thin interbeds of sand and mud (Hambrey et al., 1989). The CIROS-1 core was originally interpreted in terms of depositional setting, ice proximity and water depth (Hambrey et al., 1989; Hambrey & Barrett, 1993; Barrett, 1996). The breccia at the base of the hole is interpreted as a fault-brecciated conglomerate. The overlying sandstone/mudstone/ diamictite succession is marine, inﬂuenced to varying degrees by resedimentation and iceberg-rafting. Above the Early/Late Oligocene hiatus, the sandstones were regarded as ﬂuvial, and the diamictites as basal glacial deposits, indicating ice overriding the site. However, Fielding et al. (1997) argued, based on a sequence stratigraphic analysis, that the Late Oligocene diamictite was also glaciomarine. In contrast, Hiemstra (1999) reverted in part to the original view of grounded ice on the basis of microstructural studies. Whichever solution is the correct one, there is no clear evidence for a major environmental shift at the E/O boundary, but there is one towards lower sea level at the Early/Late Oligocene transition. A record of climate change through the E/O transition has also been determined from the environmental magnetic record in the CIROS-1 core (Sagnotti et al., 1998). Variations in magnetite were related to the concentration of detrital material transported into the Victoria Land Basin, inﬂuenced by climate and weathering rates on the Antarctic continent (especially of the Ferrar Group). Sagnotti et al. (1998) determined, from changes in the abundance of magnetite, that although there were some cold dry intervals (35–36 and W36.5 Ma) alternating with warm humid climates during the Late Eocene, a stable cold dry climate was not established in Antarctica until the E/O boundary, with major ice-sheet growth occurring at the Early/Late Oligocene boundary. This pattern matches clay mineral history, which shows a shift from smectite-rich to smectite-poor assemblages

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in Antarctica at the E/O boundary (Ehrmann and Mackensen, 1992; Ehrmann, 1997). The second core that contains the E/O transition is the Cape Roberts Project Core CRP-3; Barrett (2008; Chapter 3) has synthesized the wealth of data and numerous papers from this and other Cape Roberts cores. The strata were deposited in the same rift basin, the Victoria Land Basin, as CIROS-1. The basin ﬂoor comprises Early Devonian sandstone, above which is about 1,500 m of Cenozoic sediment. CRP-3 records a dolerite conglomerate and a basal sandstone breccia at the base of the Cenozic succession, believed to be of latest Eocene age (34 Ma). Above lies nearly 800 m of sandstone with thin beds of conglomerate, all of Early Oligocene age. Diamictite and sandstone occur towards the top of the hole, while outsized clasts are scattered through much of the core. Core CRP-2A/2A almost follows on directly above CRP-3 and spans the Late Oligocene/Early Miocene interval. The CRP-1 core overlaps with the top of CRP-2/2A core and extends further up into the Miocene. Alternating sequences of mudstone, sandstone, diamictite and conglomerate occur within this part of the Cenozoic record. The facies in CRP-3 and CRP-2A/2A represent a marine sedimentary record (Barrett, 2007). Conglomerate, sandstone and mudstone are typical of the coastal margin of a subsiding sedimentary basin, inﬂuenced by iceberg rafting. The diamictite beds additionally record tidewater glaciers that extended periodically beyond the coast. In these respects, the combined Cape Roberts cores represent an expanded record of that from CIROS-1, although there is little evidence of ice-grounding. Nevertheless, the facies ﬂuctuations are thought to reﬂect glacioeustatic changes in sea level on a wave-dominated coast, in parallel with tidewater glacier ﬂuctuations. In this context, diamicton and sand represent nearshore sedimentation, and this association grades upwards into shelf mud and then to inner shelf/shoreface sand. A conceptual depositional model for the Late Oligocene/Miocene interval was developed by Powell et al. (2000), based on facies associations and comparison with modern glaciomarine environments, such as those in Alaska and Greenland. This shows that during an advance and still-stand, a grounding-line fan develops, and this is followed by rapid recession until another fan develops. The sequence becomes even more complex when the glacier overrides previously formed fans. Figure 8.3 is a simpliﬁed version of this model. The huge volume of palaeoecological, mineralogical and geochemical data generated by these drilling projects enables us to gain insight into the broader environmental and climatic evolution through the early Cenozoic (Hambrey et al., 2002; Barrett, 2007). A temperate glacier regime is suggested for the Early Oligocene, followed by cooling into the Miocene, typiﬁed by

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Figure 8.3: Grounding-line fan model of glaciomarine sedimentation for Late Oligocene/Early Miocene time (from Hambrey et al., 2002; Simpliﬁed from Powell et al., 2000). Reproduced with permission of The Geological Society Publishing House, Bath, UK.

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Figure 8.4: Cartoon depicting the Victoria Land coast in Early/Late Oligocene time, with glacier- and river-inﬂuenced coast, and vegetated mountainsides and lowlands. This scenario is based on a combination of sedimentological evidence and ﬂoral data from CIROS-1 and Cape Roberts cores (from Hambrey et al., 2002). Reproduced with permission of The Geological Society Publishing House, Bath, UK. polythermal glaciers. The cold frigid regime of today only began at the end of Pliocene time. The Early Oligocene landscape was characterized by temperate glaciers ﬂowing from the early East Antarctic Ice Sheet, some terminating in the sea, and others on braided outwash plains (Fig. 8.4).

8.2.5. The Prydz Bay Region Drilling in Prydz Bay was undertaken by two ship-borne legs of the ODP (Fig. 8.5). In contrast to the western Ross Sea cores, core-recovery rates here were much less satisfactory; hence interpreting depositional processes is more questionable. Nevertheless, plausible scenarios have been derived, albeit lacking precise constraints owing to core loss. Prydz Bay represents the continuation of the Lambert Graben, which contains the Lambert Glacier–Amery Ice Shelf System, an ice drainage basin covering approximately 1 M km2, draining 13% of the East Antarctic Ice Sheet by surface

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Figure 8.5: Location of drill sites in Prydz Bay, East Antarctica, from ODP legs 119 and 188. Continental shelf sites 742 and 1166 include strata the cross the E/O transition. From shipboard Scientiﬁc Party (2001a). Reproduced with permission of the Ocean Drilling Program, College Station, Texas. area. Thus, the record in Prydz Bay provides a signal of the ice sheet as a whole since its inception, and complements the Oligocene (?) to Pliocene uplifted glaciomarine record in the Prince Charles Mountains (see Haywood et al., 2008; Chapter 10). Prydz Bay itself is dominated by a trough-mouth fan that prograded during phases of glacier advance to the shelf break. Like the western Ross Sea, large data-sets are available covering all aspects of core analysis from ODP Legs 119 and 188 (Barron, Larsen, & Shipboard Scientiﬁc Party, 1991; Cooper and O’Brien, 2004; Cooper et al., 2004) and a convenient summary has been provided by Whitehead et al. (2006). ODP Leg 119 obtained two cores, 739 (480 m) and 742 (316 m), the lower parts of which were loosely dated as Middle Eocene to Early Oligocene. The dominant facies recovered was massive diamictite, with minor stratiﬁed diamictite and sand (Hambrey et al., 1991, 1992). Poorly consolidated ﬁnegrained facies may well have been washed away during the drilling, however, since core-recovery rates were less than 50%. A few broken shell fragments are present, but there is a dearth of material suitable for precise dating. The base of

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Site 742 is represented by a zone of soft-sediment deformation. The Oligocene succession forms part of a prograding unit as deﬁned in seismic proﬁles, but is truncated by a regional unconformity. Above lies a ﬂat-lying sequence of more diamictite, some with preferred clast orientation and overcompaction, of late Miocene to Pliocene age (Cooper et al., 1991). Leg 188 drilled Site 1166 on the continental shelf near Sites 739 and 742 in order to obtain a more complete record of the E/O transition, but again core recovery was poor (19%). From the base upwards, Late Eocene matrixsupported sand was followed by a transgressive surface and the Late Eocene to Early Oligocene graded sand and diatom-bearing claystone with dispersed granules. These facies were capped, above an unconformity, by ‘‘clast-rich clayey sandy silt’’ (diamicton/ite) of Neogene age (Shipboard Scientiﬁc Party, 2001a,b,c,d). The interpretation of the Leg 119 facies is as follows: the deformed bed at the base of Site 742 may represent the ﬁrst stages of glaciation, with the ancestral Lambert Glacier extending across the continental shelf for the ﬁrst time. Then the bulk of the recovered facies in Sites 739 and 742 (diamictite) records deposition from the grounding-line of a tidewater glacier margin, by debris rain-out and submarine sediment gravity ﬂow beyond the shelf break, conditions which characterize much of Early Oligocene time. The overlying Miocene succession is quite different, and represents successive advances across the already prograded shelf to produce sub-glacial, ice-proximal and ice-distal facies in alternation, forming the ﬂat-lying sequence (Hambrey et al., 1991). Leg 188’s Site 1166 begins at the base with Late Eocene sand of ﬂuvio-deltaic character, and is inferred to be pre-glacial. The overlying Late Eocene to Early Oligocene sand and claystone represent shallow marine and open marine conditions, respectively, but in a proglacial setting as indicated by ice-rafted granule-sized material. The Neogene strata that lie unconformably above represent full glacial conditions. Combining the sedimentary and seismic records, along with bathymetric data from the over-deepened Lambert Graben, a conceptual model of erosion and deposition has been developed (Fig. 8.6). Considering the whole Prydz Bay region in 3D, it is apparent that Early Oligocene time was characterized by sedimentation at the shelf break. A major change then took place in the Late Miocene to mid-Pliocene following the development of a fast-ﬂowing ice stream that now occupied a more constrained channel, leading to the growth of the trough-mouth fan (Shipboard Scientiﬁc Party, 2001a,b,c,d; Taylor et al., 2004). Comparing data-sets from the western Ross Sea and Prydz Bay, it is apparent that ice ﬁrst reached the continental shelf edge in earliest Oligocene time in Prydz Bay, while the Victoria Land coast was inﬂuenced by iceberg

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Figure 8.6: Conceptual model of styles of glacial sedimentation during phases of advance and recession of the ancestral Lambert Glacier (from Hambrey and McKelvey, 2000b). (A) ‘‘Cold polar’’ glaciation as occurs today. (B) Expanded ice in the Neogene Period. (C) Recessed ice in the Neogene period. Both B and C are characterized by polythermal glaciation (comparable to the High Arctic today), and may also have been a feature of the Palaeogene period. Data are too limited, however, to deﬁne the palaeoclimate for this earlier period. Reproduced with permission of The Geological Society of America, Boulder, Colorado.

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rafting. Late Oligocene time saw repeated ice expansions recorded in the CIROS-1 core by a succession of basal tills, and probably by the major unconformity in Prydz Bay. Subsequently, frequent expansions took place to the shelf edges in both areas, resulting in a fragmented stratigraphic record.

8.3. Climate Signals from the Terrestrial Realm – Fossil Plants and Palynomorphs 8.3.1. Plant Macrofossils An important source of palaeoclimate information on land comes from fossil plants (both macro- and microfossils) (Fig. 8.7). The fossil plant record suggests that during the Late Palaeocene to Early Eocene moist, cool temperate rainforests were present, similar to modern low to mid-altitude Valdivian rainforests in southern Chile. These forests were dominated by Nothofagus and conifer trees, with ferns, horsetails and some less-prominent angiosperm groups. Assemblages of Eocene fossil plants and palynomorphs signal a generally warm climate during the Early Eocene but conditions deteriorated throughout the Antarctic Peninsula through the latter part of the Eocene when cold, seasonal climates developed (Francis, 1991; Francis and Poole, 2002; Francis et al., 2004, 2008; Poole et al., 2005). Fossil plant assemblages of younger age have been found in glacial sediments in the Transantarctic Mountains, representing vegetation that grew on Antarctica under icehouse conditions. These include the Sirius Group ﬂora from the Transantarctic Mountains (the age of which is problematic, Francis and Hill, 1996; Wilson et al., 1998a,b; Askin et al., 1999), and a new ﬂora of Miocene age discovered in the Dry Valleys (Ashworth et al., 2008). Surprisingly, leaf fossils, although single leaves, have also been discovered in drill cores within Oligocene and Miocene glacial sequences. The early glacial world of Antarctica was clearly not totally barren of vegetation. Palaeogene fossil plants have been discovered from around the Antarctic margin in outcrop, in sea-ﬂoor cores and in glacial erratic boulders from the Antarctic Peninsula (King George Island and Seymour Island) and Ross Sea (McMurdo Sound and Minna Bluff) regions. The collections consist of compressions, impressions and petrifactions of leaves, seeds, ﬂowers and wood indicating a high southern latitude ﬂora of variable diversity but dominated by fossils comparable to modern Nothofagus (the southern beech) and conifer trees. Ferns, horsetails and additional signiﬁcant Southern

PatagonianMagellanic rainforests, Chile

Valdivian rainforests, Chile

NothofagusPodocarpus dominated communities

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the small leaves suggest this may be too warm (Francis, 1999).

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Recent fern-bush communities of southern oceanic islands, e.g. Gough and Auckland Islands

Nothofagus forest with well-developed undergrowth, where ferns (including tree ferns) are important

MAT c.10°C MAP c.1000mm

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Figure 8.7: Antarctic Peninsula Late Palaeocene–Late Eocene fossil ﬂoras: composition, modern analogue vegetation and palaeoclimate interpretations. Refer to text for relevant references. MAT: mean annual temperature; MAP: mean annual precipitation.

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Stratigraphic Flora Composition age King George Island, South Shetland Islands: Cytadela/Platt Ferns Late Eocene Cliffs Nothofagus-type (mostly small) and other dicotyledon types ?Podocarpaceae Petrified Forest Fagus-Nothofagus Creek Araucaria Mount Wawel Late Equisetum Ferns PalaeoceneSeveral Nothofagus spp. Middle (microphyllous leaves) and a Eocene few other angiosperms Podocarpaceae Dragon Glacier Equisetum Moraine Ferns Angiosperms (dominated by Nothofagus spp.) Conifers (Araucariaceae, Cupressaceae, Podocarpaceae)

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Composition

40 taxa: Ferns Mixed broadleaf angiosperms (dominated by Nothofagus spp.) Conifers (podocarp, araucarian, cupressacean)

Ferns Angiosperms (including Nothofagoxylon spp., Weinmannioxylon eucryphioides, Myceugenelloxylon antarcticus Conifers (Araucaria, Cupressinoxylon, Podocarpoxylon fildesense)

Flora

Fossil Hill

Collins Glacier

3

Figure 8.7: (Continued ).

Temperate – warm temperate3 Slightly above 10 C with a low annual temperature range (MAT c.9 C), Abundant precipitation (c.1000mm) Cool temperate

Inferred palaeoclimate

Valdivian and Cool to cold temperate Magellanic rainforests, (MAT 11-13°C or Chile 10.8°C) with abundant rainfall (MAP 10003000mm) Seasonal.

Low-mid altitude Valdivian rainforests, Chile

Cool temperate rainforest

Cool – cold temperate rainforest

Tropical Latin America and southern South American rainforests

Modern analogue

Rainforest – mixed neotropical and subantarctic elements2

Vegetation model

Perhaps altitude related? The Nothofagus leaves in the Fossil Hill flora are much bigger than their modern relatives suggesting a warmer and more humid climate during the Middle Eocene.

2

Seymour Island, James Ross Basin: La Meseta Fm. Ferns, angiosperms and conifers Early-latest including: Nothofagus spp. Late Eocene (notophyllous at older Middle Eocene locality), Dilleniaceae, Myricaceae, Myrtaceae, Lauraceae, Proteaceae, Podocarpaceae, Araucariaceae

Stratigraphic age Late PalaeoceneMiddle Eocene (continued)

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Hemisphere angiosperm families, including the Proteaceae, Myrtaceae and Lauraceae, are also represented. The macrofossil record has largely been described from isolated collections, even of single leaves in cores, and although reasonably dated as a whole, a comprehensive understanding of stratigraphic relationships between the ﬂoras, particularly in the Antarctic Peninsula region, is at present hampered by differing stratigraphic interpretations. 8.3.2. Antarctic Peninsula: King George Island Many macroﬂoras have been discovered on King George Island in the South Shetland Island group, north of the Bransﬁeld Strait in the Antarctic Peninsula region, currently dated between Late Palaeocene and Late Eocene in age. The ﬂoras may have lived at a palaeolatitude of B621S, similar to its present-day location (Lawver et al., 1992). The stratigraphy is complex because Birkenmajer (1981, 1989, 1990) and Birkenmajer et al. (1986) erected many local formations in comparison to a simpler scheme created by Smellie et al. (1984). No single stratigraphic scheme exists and so the relationship between the ﬂoras is confusing. The stratigraphic framework used here includes both schemes (also reviewed by Hunt, 2001). Leaf macroﬂoras, currently understood to be of Late Palaeocene to Middle Eocene in age, have been described in varying completeness from the Admiralty Bay and Fildes Peninsula areas of the island. In the Admiralty Bay area, the Middle Eocene Mt. Wawel Formation (Point Hennequin Group) contains the macroﬂora deposits collectively known as the Point Hennequin Flora with individual localities named Mount Wawel and Dragon Glacier Moraine ﬂoras (Zastawniak et al., 1985; Birkenmajer and Zastawniak, 1989a; Askin, 1992; Hunt, 2001; Hunt and Poole, 2003). The Mount Wawel ﬂora comprises macrofossils of Equisetum (horsetail), ferns and several Nothofagus species as microphyllous leaves (a leaf-size category of 2.5–7.5 cm long), in addition to a few other angiosperms and Podocarpaceae. The Dragon Glacier Moraine ﬂora is similar, the angiosperm leaves being dominated by Nothofagus and the conifers including Araucariaceae and Cupressaceae, in addition to Podocarpaceae. The Middle Eocene Petriﬁed Forest Creek ﬂora from the Arctowski Cove Formation and the Late Eocene Cytadela ﬂora from the Point Thomas Formation are both within the Ezcurra Inlet Group. The former is a wood ﬂora requiring revision, but intermediate Fagus– Nothofagus-type species are recorded. The Cytadela leaf ﬂora includes ferns (including a Blechnum-afﬁnity species), mostly small Nothofagus-type

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leaves with pinnately veined leaves of other dicotyledonous types and possible Podocarpaceae (Birkenmajer and Zastawniak, 1989a; Askin, 1992; Birkenmajer, 1997). Birkenmajer and Zastawniak (1989a) considered this ﬂora to be E/O boundary in age. In this region, therefore, Nothfagus-dominated forests were the norm in the Middle to Late Eocene with ferns and tree ferns becoming increasingly important. Estimated mean annual temperatures of 5–81C are slightly cooler than those on Seymour Island to the east during the Middle Eocene, and the vegetation was similar to the southernmost Patagonian–Magellanic forests of southern Chile (Zastawniak et al., 1985; Birkenmajer and Zastawniak, 1989a; Askin, 1992; Hunt, 2001). By the Late Eocene, vegetation was more comparable to the recent fern-bush communities of southern oceanic islands (e.g. the Auckland Islands), interpreted from the Cytadela and Petriﬁed Forest Creek ﬂoras (Birkenmajer and Zastawniak, 1989a; Askin, 1992; Birkenmajer, 1997). However, mean annual temperature estimates of 11.7–151C appear too high, especially considering the small-sized leaves (Francis, 1999). In the Fildes Peninsula area in the southwest of King George Island, the Fildes Peninsula Group contains the contemporary Middle Eocene Collins Glacier and Rocky Cove ﬂoras within the Fildes Formation, and the diverse Late Palaeocene–Middle Eocene Fossil Hill ﬂora (Fossil Hill Formation). The latter is a leaf ﬂora containing 40 recognized taxa, including mixed broadleaf angiosperms (with large-leaved Nothofagus species), conifers (podocarp, araucarian and cupressacean) and ferns (Birkenmajer and Zastawniak, 1989a,b; Li, 1992; Haomin, 1994; Francis, 1999; Reguero et al., 2002). Neotropical and sub-Antarctic elements appear to be mixed perhaps indicating a collection derived from communities at different altitudes (Li, 1992), although this mixed signature may be a feature of early Tertiary polar vegetation (Francis et al., 2004). The Nothofagus leaves are much larger than their modern relatives, suggesting a warmer and more humid climate during the early part of the Eocene. Estimates of mean annual temperature suggest W10 1C (from 40% entire-margined leaves) and a small annual temperature range (Li, 1992). Fossil leaves remain undescribed from the Rocky Cove ﬂora; however, wood from this locality has been identiﬁed as Nothofagoxylon antarcticus (Shen, 1994; Hunt, 2001). The Collins Glacier deposit is primarily a wood ﬂora that includes wood of both coniferous (Cupressinoxylon sp. and Podocarpoxylon ﬁldesense) and angiospermous (Nothofagoxylon spp., Weinmannioxylon eucryphioides (Cunoniaceae) and Myceugenelloxylon antarcticus (Myrtaceae)) afﬁnity (Hunt, 2001; Poole et al., 2001, 2005). The mean annual temperature had dropped to 91C by the Middle Eocene and

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precipitation appears to have increased; however, the latter is probably due to changes in the environmental setting rather than climate change because of no immediate change in angiosperm wood from a semi-ring to ring porous condition (Poole et al., 2005). 8.3.3. Antarctic Peninsula: Seymour Island The Late Eocene La Meseta Formation of Seymour Island contains ﬂoras that document climate change. Leaves and wood of both angiosperm and conifer afﬁnity occur with fern fossils and a ﬂower (Case, 1988; Reguero et al., 2002; Francis et al., 2004). Middle Eocene Nothofagus leaves were found to be notophyllous (a leaf-size category 7.5–12.5 cm long). Angiosperm fossils afﬁliated to families including Nothofagaceae, Dilleniaceae, Myricaceae, Myrtaceae, Elaeocarpaceae, Moraceae, Cunoniaceae, Winteraceae and Lauraceae have been described (Gandolfo et al., 1998a,b; Reguero et al., 2002; Francis et al., 2004). Doktor et al. (1996) also described leaves afﬁliated with Podocarpaceae, Araucariaceae, Nothofagaceae and Proteaceae. Gothan (1908) was the ﬁrst to describe fossil wood from Seymour Island, which has subsequently been re-examined by several authors and identiﬁed as having both angiosperm and coniferous afﬁnities (Francis, 1991; Torres et al., 1994; Brea, 1996, 1998; Francis and Poole, 2002; Reguero et al., 2002). Decrease in leaf sizes during the Late Eocene suggests that climate deteriorated towards the end of the Eocene, as observed in studies of the La Meseta Formation by Case (1988) and Reguero et al. (2002). Gandolfo et al. (1998a,b) suggested a MAT of 11–131C for the Cucullaea I Allomember during the early Late Eocene. Further climate data were provided by leaf margin analyses of a Late Palaeocene ﬂora (Cross Valley Formation) and of the early Late Eocene Cucullaea 1 ﬂora, which indicate a decrease in ﬂoral diversity and a change from mean annual temperatures of 141C in the Late Palaeocene to 111C in the early Late Eocene, with signs of freezing winters in the Late Eocene (Francis et al., 2004). 8.3.4. Ross Sea Region: McMurdo Sound Only two E/O ﬂoras have been described from East Antarctica – a signiﬁcant, but ex situ, Middle–Late Eocene ﬂora within glacial erratic boulders found at Minna Bluff (mentioned above), and a single Nothofagus leaf within the Early Oligocene strata of the CRP-3 core in McMurdo Sound

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(Cantrill, 2001; Florindo et al., 2005). An additional single Nothofagus leaf was also found in the CIROS-1 core, originally thought to be Oligocene in age (Hill, 1989), but after recent reﬁnement of the age model, it is now considered Early Miocene (Roberts et al., 2003). The McMurdo erratic ﬂora comprises leaves, wood, seeds and Araucaria-type scale-leaves (Pole et al., 2000). Two species of Nothofagus leaves were found, some of which are interpreted as deciduous based on the interpretation of plicate vernation (venation with distinct folds). Fossil wood was identiﬁed as Araucarioxylon, Phyllocladoxylon and Nothofagoxylon (Francis, 1999, 2000). The CRP-3 Nothofagus leaf (compared to Nothofagus beardmorensis from the Sirius Group; Francis and Hill, 1996; Cantrill, 2001) was small and also had plicate vernation, but differs to the leaf in the CIROS-1 core thought to be of N. gunnii afﬁnity, an alpine species from Tasmania (Hill, 1989; Francis, 1999). The ﬂoras in the erratics indicate the presence of large forest trees comparable to those in South American araucarian (emergent) – Nothofagus forests in the Valdivian Andes of Chile or the Phyllocladus- and Nothofagusdominated cool, sclerophyll forests of temperate New Zealand and Tasmania. Francis (1999, 2000) suggested a cool temperate climate, with a mean annual temperature of o131C from the wood ﬂora, considering some winter snow likely, but temperatures not cold enough to allow extensive ice to form at sea level. Single Nothofagus leaves found in the CRP-3 and CIROS-1 cores, from Early Oligocene and Early Miocene intervals, respectively, indicate, in conjunction with the palynomorphs, a cold temperate and periglacial climate at those times. The CRP-3 Early Oligocene vegetation is compared to low Nothofagus woodland in the Magellanic region of southern Chile (Francis and Hill, 1996; Cantrill, 2001). Cantrill (2001) compared it to N. beardmorensis, known from the Sirius Group and considered to have a minimum requirement of 221C (which is probably conservative) with several weeks at least 51C during the growing season. The Early Miocene Nothofagus leaf is different and may have come from a smallto medium-sized tree in a sub-alpine rainforest or shrub community probably o1 m tall and living in exposed conditions (Hill, 1989; Francis, 1999). 8.3.5. Environmental Signals from Palynomorphs While macrofossils of plants, found usually in distinct levels in geological sections, allow a fair understanding of the biodiversity and ecology of fossil vegetation at speciﬁc times, palynological assemblages can deliver a higher resolution picture of vegetation and climate change through time, especially

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due to their presence in drill cores. In addition, pollen and spores can be recovered from areas in which macrofossils are unknown (e.g. Prydz Bay). Palynological studies have been undertaken on various Palaeogene sections in West and East Antarctica, but complete recovery is rare, especially over the crucial time interval spanning the E/O boundary when climate dramatically deteriorated. Cranwell (1959) carried out the earliest palynological studies in the Antarctic realm on a single sample of probable Palaeogene age from Seymour Island in the Antarctic Peninsula area. Subsequently, several early Tertiary stratigraphic sections on Seymour, Cockburn and King George Islands, and cores from the South Orkney Islands (ODP Leg 113, Site 696) and South Scotia Ridge (Bruce Bank, Eltanin Core IO 1578-59) have been subjected to detailed palynological analyses (Mohr, 1990; Askin et al., 1991, 1997; Grube, 2004; Grube and Mohr, 2004). Data used to reconstruct Tertiary vegetation and climate in East Antarctica have been derived from drilling campaigns in southern McMurdo Sound in the Ross Sea (CIROS-1, CRP-2/2A and CRP 3: Mildenhall, 1989; Askin and Raine, 2000; Raine and Askin, 2001; Prebble et al., 2006) and in Prydz Bay (MacPhail and Truswell, 2004). ODP Leg 189, in the Tasman Sea, cored the E/O boundary (Grube and Mohr, 2008). The glacial erratics of Eocene-aged sediments in southern McMurdo Sound have also yielded terrestrial palynomorphs. Pollen assemblages in Antarctic Palaeogene strata contain many taxa comparable to those still found today in southern high latitudes, including areas of southern South America, Tasmania, Australia, New Zealand and New Caledonia. They vary quite substantially in the abundance of their major components, and consist of moss and fern spores, gymnosperm and angiosperm pollen. Ferns were species-rich until the Middle to early Late Eocene (Mohr, 2001) and include genera that live today under humid subtropical conditions, such as Cnemidaria (Mohr and Lazarus, 1994). At Bruce Bank (c. 46–44 Ma), fern spores dominated the assemblage, at intervals comprising more than 50% of the sporomorphs. During the Late Eocene, and even more so during the Oligocene, fern diversity and abundance dropped dramatically. During the Early Miocene, some of the taxa seem to return, but recycling of older Eocene spores cannot be completely ruled out, a problem prevalent in glaciogene sediments all around the Antarctic. Gymnosperms were important components of the southern high latitude Palaeogene forests. Cycads seem to have been present until the Middle Eocene (Cycadopites), while Araucariaceae (Araucaria), Cupressaceae and especially Podocarpaceae pollen (Podocarpus, Phyllocladus, Lagarostrobus,

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Dacrydium and Microcachrys) play a major role in the palyno-associations well into the Oligocene. Angiosperms were relatively species-rich in the Palaeogene Antarctic pollen spectra with a dominance of various types of Nothofagidites (pollen comparable to that of extant Nothofagus, the southern beech), but diversity declined during the Late Eocene. Middle to Late Eocene assemblages from the Antarctic Peninsula area (ODP Leg 113, Site 696) show, in earliest sections, relatively large amounts of angiosperm pollen with a clear dominance of Nothofagidites, while in assemblages of latest Eocene and Early Oligocene age, moss spores become more common, and, except for Nothofagidites, almost no angiosperms are registered. The McMurdo Sound cores (CIROS-1, CRP-2/2A and CRP-3, Late Eocene to Oligocene) and glacial erratics (Middle to Late Eocene) are characterized by prominent Nothofagidites, particularly Nothofagidites lachlaniae and the N. fusca group. Various Podocarpus taxa are also abundant (Askin, 2000; Raine and Askin, 2001). In a relatively short sequence of CRP-3 (glaciomarine cycle 26) of Early Oligocene age, N. fusca-type, N. ﬂemingii and N. lachlaniae contribute each about 23% of the total count (Prebble et al., 2006). In glaciomarine cycle 11 of Late Oligocene age, N. fusca-type pollen dominates with about 50%, followed by N. ﬂemingii and N. lachlaniae. In Prydz Bay sections dated Late Eocene, Nothofagidites is clearly dominant at 41–57% of sporomorphs; the second largest group are conifer pollen that reach in a few samples up to 50% and more. Fern spores comprise 6–20% and cryptogams 3–6% (Macphail and Truswell, 2004) (Figs. 8.8 and 8.9). During the warmer periods of the Palaeogene, the following families have been identiﬁed from pollen: Aquifoliaceae (includes holly), Casuarinaceae (she-oak), Cunoniaceae/Elaeocarpaceae, Epacridaceae (southern heath, now included within Ericaceae), Euphorbiaceae (spurge), Gunneraceae, Liliaceae, Myrtaceae, Nothofagaceae (including all four morphotype groups brassiitype, fusca-type, menziesii-type and ancestral-type; Dettmann et al., 1990) Olacaceae, Proteaceae (Gevuina/Hicksbeachia, Adenanthos, Carnarvonia, Telopea and Beauprea; Dettmann and Jarzen, 1991), Restionaceae (rush), Sapindaceae (soapberry) and Trimeniaceae (Prebble et al., 2006). In Prydz Bay sections (Late Eocene) and within the La Meseta Formation on Seymour Island (Eocene), Fischeripollis and Droseridites, that belong to Droseraceae (sundew), which are today restricted to moors or damp sites, are excellent ecological markers. During the Late Eocene to Oligocene, Apiaceae, Asteraceae (daisy), possibly Campanulaceae (bellﬂower), Caryophyllaceae (carnation), Chenopodiaceae (now in the Amaranthaceae), Onagraceae (willowherb; Corsinipollenites) and Gramineae (grasses) seem to play a role as members of

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Figure 8.8: Relative abundance (%) of major plant groups in ODP Leg 189, Site 1168. Asterisk marks the E/O boundary. the local vegetation (Mildenhall, 1989; Mohr, 1990; Askin, 1997; Askin and Raine, 2000; Raine and Askin, 2001; Prebble et al., 2006). During the Early Oligocene, a low shrub or closed Nothofagus-podocarp forest of small stature may have developed, occupying warmer sites on the Antarctic continent (Prebble et al., 2006). In colder phases, a tundra-like vegetation, evidenced by moss spores, few but relatively diverse herb pollen and a few Nothofagidites pollen, derived possibly from dwarfed southern beech, may have grown near the coast. Palynological studies by Grube and Mohr (2008) of cores from ODP Leg 189, Site 1168 in the Tasman Sea show a clear response to E/O climate change. The abundance-time-chart (Figs. 8.8 and 8.9) for the Tasman Sea samples shows that during the latest Eocene, the pollen ﬂora was dominated by the Nothofagaceae (especially the evergreen type Brassospora), with araucarian and podocarp conifers (gymnosperms) and typical fern families (cryptogams). Near the E/O boundary itself, there is a short peak in the occurrence of araucarian and some other gymnosperm pollen, as well as an increase in ferns, in response to a decline in Nothofagaceae. Surprisingly, however, there was no sustained change in terrestrial pollen after this that might reﬂect a major change in climatic regime. Vegetation typical of latest Eocene composition seems to have been restored during the earliest Oligocene, signiﬁcant changes

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Figure 8.9: Relative abundance (%) of sporomorphs in the Tasman Sea samples from ODP Leg 189, Site 1168. Asterisk marks the E/O boundary; Lines represent palynological ‘‘events’’. being the decrease in Casuarinaceae angiosperms, and the gradual replacement of Osmundaceae ferns by Schizaeaceae and Gleicheniaceae. Nothofagus fusca- and menziesii-type pollen increase from about 33.8 Ma on, while other angiosperm pollen show a slight decline, which Grube and Mohr (2008) interpret as a gradual response to long-term cooling. The pollen record in Figs. 8.8 and 8.9 also highlights a further short-lived episode of vegetational change at about 32.9 Ma. The pattern is similar to that at 33.7 Ma, with an increase in araucarian conifer pollen and in fern spores at the expense of the angiosperms, especially the Nothofagaceae – does this represent a later episode of cold climate? The pollen diagram in Fig. 8.9 also hints at cyclical changes, possibly at intervals of 0.8 m.y or even 0.4 m.y.

8.4. Environmental Changes Documented by Marine Microfossils Early Palaeogene marine microfossil associations from the circum-Antarctic realm are typically characterized by the dominance of largely endemic

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Antarctic organic walled dinocysts and siliceous groups like diatoms and radiolarians over calcareous microfossils (e.g. Brinkhuis et al., 2003; Stickley et al., 2004; Warnaar, 2006). The characteristic endemic dinocyst assemblages reported from Middle and Late Eocene deposits around Antarctica are often referred to as the ‘‘Transantarctic Flora’’ (cf. Wrenn and Beckmann, 1982). In fact, Late Cretaceous to (early) Palaeogene organic walled dinocysts from the circum-Antarctic realm are comparatively well known, notably from southern South America, from James Ross and Seymour islands, but also from southeastern Australia and New Zealand, from erratics along the Antarctic margin, from the Ross Sea continental shelf (CIROS; CRP) and from several ocean drill sites (see, e.g. Haskell and Wilson, 1975; Wilson, 1985, 1988; Askin, 1988a,b; Wrenn and Hart, 1988; papers in Duane et al., 1992; Pirrie et al., 1992; Mao and Mohr, 1995; Hannah, 1997; Truswell, 1997; Hannah et al., 2000; Levy and Harwood, 2000a,b; Guerstein et al., 2002). Meaningful chronostratigraphic calibration of (sub-)Antarctic dinocyst events was a classic problem due to the general absence of other ageindicative biotas and/or magnetostratigraphy or other means of dating in sections in which dinocysts are encountered. The ﬁrst integrated Oligocene to earliest Miocene biomagnetostratigraphy, including dinocysts, was achieved only relatively recently on the basis of successions drilled during the CRP (e.g. Hannah et al., 1998, 2000). Even more recently, the ﬁrst magnetostratigraphically calibrated Late Maastrichtian to earliest Oligocene dinocyst succession was established on the basis of records drilled during ODP Leg 189, offshore Tasmania (e.g. Brinkhuis et al., 2003; Sluijs et al., 2003; Huber et al., 2004; Stickley et al., 2004). Building on these studies, Warnaar (2006) produced higher resolution records for the ODP 189 holes, and (re)analysed critical intervals from other circum-Antarctic sites like 696, 739, 1090 and 1166. In the following section, the dinocyst record for the Eocene and Oligocene is documented and the implications for our understanding of palaeoceanography at this time are discussed. 8.4.1. Palaeocene–Middle Eocene Dinocysts Circum-Antarctic early Palaeogene dinocyst associations were recovered from the Tasmanian Gateway (between Australia and East Antarctica, e.g. Brinkhuis et al., 2003), New Zealand (e.g. Wilson, 1978, 1984, 1988; Willumsen, 2000; Crouch, 2001) and Seymour Island (e.g. Askin, 1988a,b; Elliot et al., 1994). The associations are characterized by nearly identical

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composition and stratigraphic succession. While circum-Antarctic endemic taxa were present at least since the Maastrichtian (e.g. Riding and Crame, 2002; Brinkhuis et al., 2003), the taxa often referred as the ‘‘Transantarctic Flora’’ established itself in the early Palaeogene in an otherwise largely cosmopolitan assemblage. Since the late Early Eocene, the inﬂuence of the ‘‘Transantarctic Flora’’ (constituted by species such as Deﬂandrea antarctica, Octodinium askiniae, Enneadocysta partridgei, Vozzhennikovia spp., Spinidinium macmurdoense and Arachnodinium antarcticum) increases until the middle Late Eocene. While the younger part of the circum-Antarctic Middle Eocene is comparatively less well studied, with records only from New Zealand (e.g. Wilson, 1988; Strong et al., 1995), DSDP Sites 280 and 281 (Crouch and Hollis, 1996), Seymour Island (Wrenn and Hart, 1988), the Scotia Sea (Mao and Mohr, 1995), ODP Site 1172 (Brinkhuis et al., 2003) and southern Argentinean successions (e.g. Guerstein et al., 2002), a broad similarity is apparent. The younger part of the Southern Ocean Middle Eocene appears to be characterized by several important last occurrences (LOs), including those of Membranophoridium perforatum, Hystrichosphaeridium truswelliae, Hystrichokolpoma spinosum and Hystrichokolpoma truncatum (cf. Wilson, 1988; see Brinkhuis et al., 2003). 8.4.2. Late Eocene–Early Oligocene Dinocysts At the Tasmanian Gateway, the early Late Eocene dinocyst distribution forms a continuation of the Middle Eocene pattern. ‘‘Transantarctic Flora’’ species predominate, and ﬁnal acmes of Enneadocysta partridgei, the Deﬂandrea antarctica group and Spinidinium macmurdoense are recorded (Fig. 8.10). Important ﬁrst occurrences (FOs) in this phase include those of Schematophora speciosa, Aireiana verrucosa, Hemiplacophora semilunifera and Stoveracysta ornata. Towards the middle Late Eocene, FOs of Achomosphaera alcicornu, Reticulatosphaera actinocoronata and Alterbidinium distinctum and the LO of S. speciosa appear important for interregional correlation, as is the FO of Stoveracysta kakanuiensis. Vozzhennikovia spp. continues to be a common constituent of the associations (Sluijs et al., 2003). Typically, sediments representing the E/O transition are barren of organic microfossils in all ODP Leg 189 records; dinocysts brieﬂy reappear in the Early Oligocene (assigned to Chron C11-1r; Stickley et al., 2004). In this single productive sample thus far from the Early Oligocene, virtually all Transantarctic Palaeogene dinocysts have disappeared (only a single, poorly preserved, probably reworked specimen of E. partridgei was recovered;

Figure 8.10: Circum-Antarctic geographical distribution maps (Late Paleocene–Miocene) showing dinocyst endemism. Maps derived from the Ocean Drilling Stratigraphic Network (ODSN). Black areas indicate (continental) blocks that are mostly sub-aerial. Note that several blocks shown in black were partly submerged (e.g., the Ross Sea, the southern Australian margin and parts of Argentina). Shaded areas indicate mostly submerged (continental) blocks (e.g., Brown et al., 2006).

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Brinkhuis et al., 2003). The association in this sample is characterized by the abundance of taxa more typical for Tethyan waters, including an occurrence of Hystrichokolpoma sp. cf. Homotryblium oceanicum (e.g. Brinkhuis and Bifﬁ, 1993; Wilpshaar et al., 1996; Brinkhuis et al., 2003). Some of the Late Eocene dinocyst events have previously been reported from the South Australian margin, e.g. from the Browns Creek section (Cookson and Eisenack, 1965; Stover, 1975). For example, the ranges of S. speciosa, A. verrucosa, H. semilunifera and S. ornata appear useful for regional and even global correlation. Many of the ‘‘Browns Creek’’ Late Eocene dinocysts have been recorded from locations around the world, also in otherwise well-calibrated sections in central and northern Italy, including the Priabonian Type Section (Brinkhuis and Bifﬁ, 1993; Brinkhuis, 1994). It appears that these index species have slightly earlier LOs in this region than they have in Italy (Tethyan Ocean), if the records of Cookson and Eisenack (1965) and Stover (1975) are combined with more recent nannoplankton and magnetostratigraphic studies from the same section (Shaﬁk and Idnurm, 1997). This aspect may be related to the progressive global cooling during the latest Eocene (Fig. 8.11). In the rare sediments covering the E/O transition, more speciﬁcally the Oi-1 event, from the Weddell Sea and near the Drake Passage (between South America and the Antarctic Peninsula), dinocysts are typically not preserved (Gradstein et al., 2004). However, in the oldest Oligocene sediments bearing dinocysts, the Transantarctic Palaeogene dinocysts (dominant in the latest recovered Eocene sediments) are replaced by cosmopolitan taxa (Gradstein et al., 2004). This suggests that the changes in dinocyst associations in this area were at least broadly similar compared to those in the Tasman Sector. In contrast, it seems that sediments covering the E/O transition are preserved at Prydz Bay (ODP Site 739). Here a gradual change is observed from the typical Transantarctic Palaeogene dinocysts to the taxa typically found in (post-) Oligocene near-Antarctic records (Warnaar, 2006). However, the (cosmopolitan) taxa that are useful for correlation, as mentioned above, are apparently not present in the Prydz Bay records. 8.4.3. Palaeoceanography During the early Palaeogene, Antarctica was less glacierized than it is now, or not glacierized at all (e.g. Zachos et al., 2001; Pagani et al., 2005; see above). South America and Australia were still not fully separated from the Antarctic continent (e.g. Livermore et al., 2005; Brown et al., 2006), which

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Figure 8.11: Generalized (Middle Eocene) dinocyst geographical distribution map overlain with the ocean circulation pattern inferred from GCM results. Maps derived from the Ocean Drilling Stratigraphic Network (ODSN). Shaded areas indicate mostly submerged (continental) blocks (e.g., Brown et al., 2006). Abbreviations: TA-SW, Trans-Antarctic Seaway (hypothetical; see Wrenn and Beckmann, 1982); TSA-SW, Trans-South American Seaway (hypothetical; see Kohn et al., 2004); EAC, East Australian Current; p-LC, proto-Leeuwin Current; p-RG, proto-Ross Gyre; TC, Tasman Current.

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prevented the development of a (proto-) Antarctic Circumpolar Current (ACC). It was hypothesized that during the early Palaeogene, warm ocean currents from lower latitudes could reach and warm Antarctica. The opening and subsequent deepening of critical conduits (i.e. Drake Passage and the Tasmanian Gateway) towards the end of the Eocene have long been thought to have played a central role in ACC establishment and Antarctic cooling (e.g. Kennett et al., 1975; Kennett, 1977, 1978; Murphy and Kennett, 1975). However, recent advances through ocean drilling (e.g. ODP Leg 189) and coupled Global Circulation Model (GCM) experiments suggest that Eocene Southern Ocean surface circulation patterns were fundamentally different than previously thought, and that the opening and deepening of oceanic gateways were of little climatic consequence (Sloan and Huber, 2001; Huber et al., 2004; Warnaar, 2006). Instead, it is nowadays argued that the ‘‘greenhouse–icehouse’’ transition was caused by changes in greenhouse gas concentrations, rather than oceanographic changes (e.g. DeConto and Pollard, 2003a,b; Huber et al., 2004; Pagani et al., 2005). The high degree of endemism in the circum-Antarctic marine microfossil associations denies the existence of a southward-bound, warm, proto-East Australian Current as proposed by Kennett et al. (1975) and Exon et al. (2004), according to Huber et al. (2004) (Fig. 8.11). For example, a colder northward ﬂowing western boundary current, designated the ‘‘Tasman Current’’ (see Huber et al., 2004) existed off southeast Australia. GCM experiments indicate that the Eocene Southern Ocean, including the southern Paciﬁc, was dominated by clockwise gyres (Sloan and Huber, 2001; Huber et al., 2004). Moreover, several studies show that the Tasmanian Gateway had already been open to neritic water depths (i.e. o200 m) since at least the Middle Eocene (Stickley et al., 2004). Deepening to bathyal water depths (i.e. 200–4,000 m) occurred during the early Late Eocene (B35.5 Ma). The Drake Passage had possibly been open to (upper) bathyal water depths by the Middle Eocene (Eagles et al., 2006; Scher and Martin, 2006). Both tectonic events thus seem to have occurred too early to be related to the Antarctic glaciation in the Early Oligocene (i.e. the Oi-1 stable-isotope event (33.3 Ma), e.g. Miller et al., 1998; Zachos et al., 2001). Given the similar continent–ocean conﬁguration, a corollary of the GCM experiments is that the Palaeogene circum-Antarctic surface circulation should not have been fundamentally different from the Cretaceous situation. If this were the case, and as long as substantial equator–pole temperature gradients existed, then it may be expected that, throughout the Late Cretaceous to early Palaeogene time interval, circum-Antarctic waters were consistently dominated by endemic biota, particularly in environments inﬂuenced by the proposed western boundary currents. Both hypotheses

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were recently tested by Warnaar (2006) by mapping distribution patterns of circum-Antarctic dinocysts through Palaeogene times and comparing them with coupled GCM results. Warnaar (2006) conceived a model termed the ‘‘refrigerator trap’’ wherein it is hypothesized that cosmopolitan and endemic dinoﬂagellates were taken through the cold and darkness along the Antarctic continent, transported by the proto-Ross Gyre. Conceivably, taxa normally living in warmer waters (e.g. the East Australian Current) that were trapped in the gyre were unable to survive such conditions. It is conceivable that the endemic taxa (notably taxa of the ‘‘Transantarctic Flora’’ and bi-polar Phthanoperidinium echinatum group) were speciﬁcally adapted to tolerate cold conditions (41C to the freezing point), prolonged darkness and possibly seasonal sea ice.

8.5. Evolution of Ocean Temperatures and Global Ice Volume During the Eocene to Oligocene from the Ocean Isotope Record The evolution of climate during the Eocene and Oligocene can be determined from the deep-sea isotope and trace element records of ocean temperatures and ice volume. Earlier isotope work suggests that the primary transition from greenhouse to icehouse world took place during the Late Eocene and Early Oligocene, with large, permanent ice sheets appearing on Antarctica at 34 Ma (Zachos et al., 1992, 1996, 2001; Miller et al., 1998; Coxall et al., 2005). This transition was preceded by a period of long-term cooling which initiated near the Early–Middle Eocene boundary, roughly 50 Ma, following a sustained period of Early Eocene warmth. The Eocene cooling trend was not monotonic, but followed a somewhat step-like pattern with several reversals, the most substantial of which was the Middle Eocene climatic optimum (MECO) (Bohaty and Zachos, 2003; Jovane et al., 2007). By the Late Eocene, the climate on Antarctica appears to have cooled sufﬁciently to allow for the formation of small, ephemeral ice sheets, a state that persisted until B34 Ma, when most of East Antarctica became glaciated by a large ice sheet (Fig. 8.12). From that time forward, the ice sheet was a permanent feature of Antarctica. For the remainder of the Oligocene, this ice sheet waxed and waned, most likely in response to orbital forcing (Naish et al., 2001). The long-term cooling trend that facilitated the formation of continental ice sheets has been attributed to either changes in palaeogeography or the concentration of greenhouse gases. Geographical isolation of

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recently (Pagani et al., 2005) lacked sufﬁcient resolution to fully test this possibility. Recent investigations of marine cores have largely focused on improving two aspects of the Eocene and Oligocene climate reconstructions: (1) the resolution of proxy records of ocean temperature and ice volume, which has promoted the development of high-resolution, orbitally tuned records, and (2) quantifying changes in ice volume, which has spurred the development and application of palaeotemperature proxies. These studies have beneﬁted in part from efforts of the ODP to recover highly expanded, stratigraphically intact sediment sequences spanning the E/O. The latest high-resolution records show that ice volume and ocean temperatures varied in a periodic fashion with power concentrated in the long eccentricity and obliquity bands (Fig. 8.13; Coxall et al., 2005). The latter is consistent with the presence of a large polar ice sheet. Prior to the Late Eocene, however, power in the obliquity band is relatively weak (Palike et al., 2001), suggesting little to no ice volume on Antarctica. These records have also revealed distinct climatic variability coherent with the lower frequency components of obliquity with periods of 1.25 m.y. The recent development of seawater temperature proxies, Mg/Ca and TEX86 (Schouten et al., 2003), has improved estimates of ice volume from oxygen isotope records. In particular, the ﬁrst low-resolution benthic Mg/Ca records suggest that much of the d18O increase just after the E–O boundary (33.4 Ma) was the result of a substantial increase in ice volume as deep-sea temperature was fairly constant (Lear et al., 2000; Billups and Schrag, 2002). Most recently, Lear et al. (2008) have interpreted a marine cooling of B2.51C associated with this ice growth from exceptionally well preserved foraminifera well above the calcite compensation depth. Though still controversial, the magnitude of ice-volume increase would have exceeded that of the present-day Antarctic Ice Sheet.

8.6. Connection of CO2 and Ice-Sheet Inception at the E/O Boundary – Computer Modelling While the onset of major, continental-scale glaciation in the earliest Oligocene has long been attributed to the opening of Southern Ocean gateways (Kennett and Shackleton, 1976; Kennett, 1977; Robert et al., 2001), recent modelling studies suggest declining atmospheric CO2 was the most important factor in Antarctic cooling and glaciation.

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As the Drake Passage and Tasmanian Gateway widened and deepened during the late Palaeogene and early Neogene (Lawver and Gahagan, 1998), the ACC and Polar Frontal Zone (APFZ) presumably cooled the Southern Ocean by limiting the advection of warm subtropical surface waters into high latitudes (Kennett, 1977). While the opening of the Tasmanian Gateway does broadly coincide with the earliest Oligocene glaciation event (Oi-1) (Stickley et al., 2004), the tectonic history of the Scotia Sea remains equivocal. Estimates for the timing of the opening of Drake Passage range between 40 and 20 Ma (Barker and Burrell, 1977; Livermore et al., 2004; Scher and Martin, 2006), clouding the direct ‘‘cause and effect’’ relationship between the gateways and glaciation. A number of ocean modelling studies have shown that the opening of both the Drake Passage and Tasmanian gateways reduces poleward heat convergence in the Southern Ocean and cools sea surface temperatures by up to several degrees (Mikolajewicz et al., 1993; Nong et al., 2000; Toggweiler and Bjornsson, 2000). More recent, coupled atmosphere–ocean GCM simulations suggest a more modest effect, however. Huber et al. (2004) showed that the Tasmanian Gateway likely had a minimal effect on oceanic heat convergence and sea surface temperatures around the continent, because the warm East Australia Current does not travel any further south if the gateway is open or closed. The gateway’s effect on East Antarctic climate and snowfall was also shown to be minimal, pointing to some other forcing (perhaps decreasing atmospheric CO2 concentrations) as the primary cause of Antarctic cooling and glaciation. The recent development of coupled climate–ice sheet models capable of running long (W106 years), time-continuous simulations of speciﬁc climate events and transitions (DeConto and Pollard, 2003a) has allowed simulations of the Oi-1 event that account for decreasing CO2 concentrations, orbital variability and prescribed changes in ocean transport (DeConto and Pollard, 2003b; Pollard and DeConto, 2005). These simulations support the conclusions of Huber et al. (2004) as to the likely importance of CO2 by showing that, even if signiﬁcant, tectonically forced changes in ocean circulation and heat transport had occurred around the E/O boundary, they would have had only a small effect on temperature and glacial mass balance in the Antarctic interior. Therefore, Southern Ocean gateways could only have triggered glaciation if the climate system was already close to a glaciation threshold. Considering the sensitivity of polar climate to the range of CO2 concentrations likely to have existed over the Palaeogene–Neogene (Pagani et al., 2005), CO2 likely played a fundamental role in controlling Antarctica’s climatic and glacial sensitivity to a wide range of forcing mechanisms. This conclusion is supported by a number of additional

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modelling studies exploring the role of orbital variability (DeConto and Pollard, 2003b), mountain uplift in the continental interior (DeConto and Pollard, 2003a), geothermal heat ﬂux (Pollard et al., 2005), Antarctic vegetation dynamics (Thorn and DeConto, 2006) and Southern Ocean sea ice (DeConto et al., in press) in the E/O climatic transition. The results of these studies can be summarized as follows. The timing of glaciation on East Antarctica was shown to be sensitive to orbital forcing, mountain uplift and continental vegetation, but only within a very narrow range of atmospheric CO2 concentrations around 2.8 times modern levels, close to the model’s glaciation threshold. Once the glaciation threshold is approached, astronomical forcing can trigger sudden glaciation through non-linear height/mass balance and albedo feedbacks that result in the growth of a continental-scale ice sheet within 100 kyrs (Fig. 8.14). The timing of glaciation appears to be insensitive to both expanding concentrations of seasonal sea ice and changes in geothermal heat ﬂux under the continent; however, a doubling of the background geothermal heat ﬂux (from 40 to 80 mW m 2) does have a signiﬁcant effect on the area under the ice sheet at the pressure-melt point (where liquid water is present), which may have had some inﬂuence on the distribution and development of sub-glacial lakes and subsequent ice-sheet behaviour. While these modelling studies have certainly improved our understanding of the importance of atmospheric CO2 concentrations relative to other Cenozoic forcing factors, several important model-data inconsistencies remain unresolved. For example, long, time-continuous GCM-ice-sheet simulations of an increasing CO2 (warming) scenario show strong hysteresis once a continental ice sheet has formed (Pollard and DeConto, 2005). In these simulations, orbital forcing alone is not sufﬁcient to produce the range of Palaeogene–Neogene ice-sheet variability (B50–120% modern Antarctic ice volumes) inferred by marine oxygen isotope records and sequence stratigraphic reconstructions of eustasy (Zachos et al., 2001; Pekar and DeConto, 2006; Pekar et al., 2006), pointing to the importance of additional feedbacks (possibly related to the marine carbon cycle and atmospheric CO2) in controlling Cenozoic ice-sheet variability. Furthermore, several recent isotopic analyses of deep-sea cores imply ice volumes during the peak Oligocene and Miocene glacial intervals that are too big to be accommodated by East Antarctica alone (Lear et al., 2004; Coxall et al., 2005; Holbourn et al., 2005). This suggests that either our interpretations of the proxy data are faulty, or episodic, bipolar glaciation occurred much earlier than currently accepted. These, among other unresolved controversies related to the climatic and glacial evolution of the high southern latitudes, will be the focus of future ACE modelling exercises and model-data comparisons.

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8.7. Summary Although no one single source of evidence or locality yields the complete story of climate change and the onset of glaciation in Antarctica, piecing together information from a range of sources, as presented above, does provide a picture of how climate cooled and glaciation became established on the continent. A summary is presented below (for references, see text above). Paradoxically, the present ice sheet hides its own history, certainly at the highest latitudes in the middle of the continent, so most of what is known about past environments is from the lower latitude marginal sites.

8.7.1. Early–Middle Eocene Polar Warmth Evidence from fossil plants, sediments and isotopes indicates that the Late Palaeocene and Early Eocene experienced warm climates at high latitudes, at least on the margins of Antarctica where strata of this age crop out. Climates appear to have been warm and wet, seas were warm and plants ﬂourished in a frost-free environment. The oldest record of glacial activity (if the dating is correct in this problematic region) is of valley-type tillites of Middle Eocene age on King George Island, indicating the presence of alpine glaciers. However, ﬂoras of Middle Eocene age from King George and Seymour islands suggest warm to cool temperate climates, generally moist and probably frost-free. The ocean isotope record also suggests that climates were generally warm until the Middle Eocene, although the climate trend was towards cooling.

8.7.2. Late Eocene Cooling A variety of sources, particularly fossil plants, suggest that during the early Late Eocene, climates cooled but perhaps not to the extent of signiﬁcant ice build-up. The Late Eocene sediment record in the Ross Sea region (McMurdo Erratics, magnetic and clay mineral record) and in the Prydz Bay area could be indicative of cold climates but the coastal/open marine shelf and ﬂuvial-deltaic environments in these two areas, respectively, do not show signs of the presence of signiﬁcant ice. By the latest Eocene, however, glacial deposits are apparent. Glacial deposits on Seymour Island, close to the E/O boundary, may indicate the presence of valley glaciers in that region, situated at about 651 South Palaeolatitude.

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The ocean record, especially marine microfossils, provides information about the climate and currents in the oceans at critical times during this interval. South America and Australia were still not separated from Antarctica during the Early Eocene, so the Antarctic Circumpolar Current was unable to develop. Instead, warm equatorial currents may have fed warmth to the continent. Palaeoceanographic changes related to the deepening of the Tasmanian Gateway and the opening of the Drake Passage are still debated but do not seem to have been strictly related to Early Oligocene climate cooling and intensiﬁcation of glaciation, so these oceanographic changes are not considered to be major drivers of polar cooling. Recent atmosphere–ocean modelling has also shown that changes in oceanography related to tectonic events were not likely to have driven the climate cooling that led to glaciation. 8.7.3. Latest Eocene/Earliest Oligocene Glaciation By the E/O boundary times, there is no doubt that ice was present on Antarctica. In the Ross Sea region, drill cores show evidence of relatively uniform marine sedimentation through the latest part of the Eocene and into the Oligocene but sediments include exotic clasts indicative of iceberg rafting. There does not appear to be a major environmental shift at this time but more of an intensiﬁcation of cooling. In the Prydz Bay region, tidewater glaciers were present in the Early Oligocene, with ice reaching the continental shelf edge. In the oceans, the oxygen isotope record and other geochemical indicators signal a strong cooling at the boundary, the Oi-1 event, which has been interpreted as a time of major build-up of ice. Even though climates were cold, vegetation was able to persist but by this time the higher diversity and warmth-loving plants of the Early and Middle Eocene forests had disappeared, to be replaced by vegetation that was dominated by several species of the southern beech, Nothofagus. Along with mosses, a few ferns and some podocarp conifers, southern beech trees probably grew as shrubby vegetation in the most hospitable areas. 8.7.4. Oligocene Ice Sheets A hiatus at the Early/Late Oligocene boundary marks a change to ﬂuvial conditions, with grounded ice or possibly glaciomarine conditions. A distinct drop in sea level is noted at this time. Facies changes and diamictite beds in the Cape Roberts core are indicative of the periodic expansion of tidewater

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glaciers, typical of a temperate glacier regime with glaciers ﬂowing from the young East Antarctic Ice Sheet across the continental margin. Throughout the Oligocene, glaciation seems to have waxed and waned. A distinct record of glacial activity is recorded by glacial sediments in the Polonez Cove Formation on King George Island of mid-Oligocene age (Krakowiak Glacial Member). At that time, ice was grounded on a shallow marine shelf. More extensive ice sheets may have been present further south in the Weddell Sea region, from which clasts of rock from the Transantarctic Mountains may have been derived to be incorporated as exotic clasts in the Polonez Cove Formation. Sediments without a glacial signature that overlie these glacial deposits suggest a phase of climate warming and glacial retreat until the next glacial pulse in the Miocene. Why did the climate cool during the Eocene and Oligocene, causing such a major change in Antarctic environments? The inﬂuence of palaeoceanographic changes is now considered less critical; instead, coupled climate–ice sheet modelling indicates that it was changing levels of atmospheric CO2 that controlled Antarctica’s climate. Factors such as mountain uplift, vegetation changes and orbital forcing all played a part in cooling the polar climate, but only when CO2 levels fell to critical threshold levels (2.8 times present-day levels) did orbital forcing tip Antarctica into its icy glacial world.

ACKNOWLEDGEMENTS Brinkhuis and Warnaar would like to thank Catherine Stickley, Matt Huber and Appy Sluijs. Francis would like to thank ACE and SCAR for funds to attend ACE meetings. B. Mohr wishes to thank for funds provided by the German Funding agency (DFG), which enabled R. Grube to study Antarctic material from various sources, mainly from Ocean Drilling Program (ODP) cores.

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Askin, R. A. (1988b). The palynological record across the Cretaceous/Tertiary transition on Seymour Island, Antarctica. In: R. M. Feldmann, & M. O. Woodburne (Eds). Geology and Paleontology of Seymour Island, Antarctic Peninsula. Geological Society of America Memoir, Vol. 169, pp. 155–162. Askin, R. A. (1992). Late Cretaceous–Early Tertiary Antarctic outcrop evidence for past vegetation and climate. In: J. P. Kennett, & D. A. Warnke (Eds). The Antarctic Paleoenvironment: A Perspective on Global Change. Antarctic Research Series. American Geophysical Union, Washington, DC, Vol. 56, pp. 61–75. Askin, R. A. (1997). Eocene to ?Earliest Oligocene terrestrial palynology of Seymour Island. In: C. A. Ricci (Ed.). The Antarctic Region: Geological Evolution and Processes. Terra Antartica Publication, Siena, pp. 993–996. Askin, R. A. (2000). Spores and pollen from the McMurdo Sound erratics, Antarctica. In: J. D. Stilwell, & R. M. Feldmann (Eds). Paleobiology and Paleoenvironments of Eocene rocks, McMurdo Sound, East Antarctica. Antarctic Research Series. American Geophysical Union, Washington, DC, Vol. 76, pp. 161–181. Askin, R. A., Ashworth, A. C., Harwood, D. H., Webb, P. N., & Wilson, G. S. (1999). Palynological survey of Sirius Group and related Neogene sediments in the transantarctic mountains. In: D. N. B. Skinner (Ed.). 8th International Symposium on Antarctic Earth Sciences, Victoria University of Wellington, New Zealand, p. 30, Programme and Abstracts. Askin, R. A., & Raine, J. I. (2000). Oligocene and Early Miocene Terrestrial palynology of the Cape Roberts Drill Hole CRP-2/2A, Victoria Land Basin, Antarctica. Terra Antartica, 7(4), 493–501. Barker, P. F., & Burrell, J. (1977). The opening of Drake Passage. Mar. Geol., 25, 15–34. Barrett, P. (2008). A history of Antarctic Cenozoic glaciation – View from the continent. In: F. Florindo, & M. Siegert (Eds). Antarctic Climate Evolution. Elsevier, Amsterdam, Vol. 8, pp. 33–83. Barrett, P. J. (Ed.) (1989). Antarctic Cenozoic History from the CIROS-1 Drillhole, McMurdo Sound. New Zealand DSIR Bulletin. Wellington, Vol. 245, 251 pp. Barrett, P. J. (1996). Antarctic paleoenvironments through Cenozoic time – A review. Terra Antartica, 3, 103–119. Barrett, P. J. (2007). Cenozoic climate and sea level history from glaciomarine strata off the Victoria Land coast, Cape Roberts Project, Antarctica. In: M. J. Hambrey, P. Christoffersen, N. F. Glasser, & B. Hubbard (Eds). Glacial Sedimentary Processes and Products. International Association of Sedimentologists, Special Publication, Vol. 39, pp. 259–287. Barrett, P. J., Hambrey, M. J., Harwood, D. M., Pyne, A. R., & Webb, P.-N. (1989). Synthesis. In: P. J. Barrett (Ed.). Antarctic Cenozoic History from the CIROS-1 Drillhole, McMurdo Sound. New Zealand DSIR Bulletin. Wellington, Vol. 245, 241–251. Barrett, P. J., Hambrey, M. J., & Robinson, P. H. (1991). Cenozoic glacial and tectonic history from CIROS-1, McMurdo Sound. In: M. R. A. Thomson,

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Developments in Earth & Environmental Sciences, 8 F. Florindo and M. Siegert (Editors) r 2009 Elsevier B.V. All rights reserved DOI 10.1016/S1571-9197(08)00009-8

Chapter 9

The Oligocene–Miocene Boundary – Antarctic Climate Response to Orbital Forcing G. S. Wilson1,, S. F. Pekar2, T. R. Naish3, S. Passchier4 and R. DeConto5 1

Department of Geology, University of Otago, P.O. Box 56, Dunedin, New Zealand School of Earth and Environmental Sciences, Queens College, CUNY, 65-30 Kissena Blvd., Flushing, NY 11367, USA and Lamont Doherty Earth Observatory of Columbia University, Palisades, NY 10964, USA 3 Antarctic Research Centre, Victoria University, P.O. Box 600, Wellington, New Zealand and GNS Science, P.O. Box 30368, Lower Hutt, New Zealand 4 Department of Earth and Environmental Studies, Montclair State University, 252 Mallory Hall, 1 Normal Avenue, Montclair, NJ 07403, USA 5 Department of Geosciences, University of Massachusetts, 233 Morril Science Center, 611 North Pleasant Street, Amherst, MA 01003-9297, USA 2

ABSTRACT Recent high-resolution Oligocene–Miocene oxygen isotopic records revealed a relatively transient, ca. 2 myr period, 1 m amplitude cyclicity in isotopic values (Oi and Mi events, respectively). Intriguingly, it has been suggested that these isotopic excursions in oceanic d18O were linked to ephemeral growth and decay in Antarctic Ice Sheets. A great deal of effort in the palaeoceanography community has been focused on developing techniques and gathering additional records to determine if the Antarctic Ice Sheet has behaved in such a transient manner in the past and indeed what factors might have led to the rapid growth and decay of ice sheets. Deciphering between temperature and ice-volume inﬂuences in the deep-sea isotopic record has proven somewhat difﬁcult. Approaches have included the sampling of sediment from beneath different water masses, development of an independent palaeothermometer using magnesium/calcium ratios and improving the resolution and accuracy of coastal sea-level records. Despite these advances, it Corresponding author. Tel.: þ64 3 4797509; Fax: þ64 3 4797527;

E-mail: [email protected] (G.S. Wilson).

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is only through the recovery of Antarctic drill core records that we have been able to test the resulting hypotheses. Combined with numerical climate models, ice-volume estimates are also available. The Antarctic Oligocene–Miocene record is most complete in the Victoria Land Basin as recovered in the CIROS-1 and CRP-2A drill cores. The strata recovered in both drill cores are cyclic in nature and interpreted to represent periodic advance and retreat of ice across the Antarctic margin in the western Ross Sea concomitant with sea-level fall and rise, respectively. Environmental data suggest a signiﬁcant Antarctic climate threshold across the Oligocene–Miocene boundary with cooler temperatures implied in the early Miocene with ice volume and palaeo-sea-level estimates suggesting a signiﬁcant but transient growth in the Antarctic Ice Sheet to B25% larger than present. The Antarctic data are entirely consistent with the predictions from deepsea records including the suggestion that the glacial advance was relatively short lived and interglacial conditions were re-established within a few hundred thousand years. The duration and transience of the Mi1 glacial expansion and swift recovery in Antarctica likely resulted from the limited polar summer warmth from the coincidence of low eccentricity and low-amplitude variability in obliquity of the Earth’s orbit at the Oligocene–Miocene boundary. This was followed by warmer polar summers and increased melt from increased eccentricity and high-amplitude variability in obliquity in the early Miocene, allowing the recovery of vegetation on the craton. Atmospheric CO2 concentrations remained below a 2 pre-industrial threshold, which promoted sensitivity of the climate system to orbital forcing. While climate and ice-sheet modelling support the fundamental role of greenhouse gas forcing as a likely cause of events like Mi1, the models underestimate the range of orbitally paced ice-sheet variability recognised in early Miocene isotope and sea-level records unless accompanied by signiﬁcant ﬂuctuations in greenhouse gas concentrations. While tectonic inﬂuences may have been secondary, they may well have contributed to oceanic cooling recorded at the Cape Roberts Project site in the South Western Ross Sea.

9.1. Introduction The paucity of Cenozoic outcrop on the Antarctic craton has led to the reliance on proxy records (isotopic signatures in microfossils, deep-sea erosion events and former sea levels on distal continental margins) to help unravel the history of climate and ice sheets on Antarctica (Kennett and Shackleton, 1976; Kennett, 1977; Wright and Miller, 1993; Miller and Mountain, 1996; Zachos et al., 2001a; Miller et al., 2005). Much attention has been focused on the search for Antarctic data sets to ‘ground truth’ signiﬁcant climate trends, events and thresholds observed in these proxy

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records. Drilling around the Antarctic margin by seven legs of the Deep Sea Drilling Project (DSDP; Kennett et al., 1974; Hayes and Frakes, 1975) and Ocean Drilling Programme (ODP; Barker et al., 1988a, b; Ciesielski et al., 1988; Barron et al., 1989; Barker et al., 1999; O’Brien et al., 2001), and several sea-ice-based drilling projects (Barrett, 1986, 1989; Cape Roberts Science Team, 1998, 1999, 2000; Fig. 9.1) has recovered Cenozoic sequences, which have allowed the testing of interpretations of Antarctic Glacial history from proxy records and climate models. While the DSDP and ODP core recovery has been between 14 and 40%, riser drilling from sea ice in the South Western Ross Sea has enabled recovered of some high-quality intervals of the Cenozoic (95–98%) recovering prima facie documentation of climate and cryospheric changes in Antarctica. One interval that is particularly well-sampled and well-dated in several drill cores is the Oligocene–Miocene boundary, permitting an accurate comparison to deepsea high-resolution isotopic records from lower latitudes (Naish et al., 2001, 2008; Wilson et al., 2002; Roberts et al., 2003). This chapter reviews recent evidence for a glacial expansion in Antarctica coincident with the Oligocene–Miocene boundary and the Mi1 deep-sea oxygen isotope excursion. The climatic signiﬁcance of the boundary has only recently become apparent from recalibration of the Oligocene–Miocene time scale using astrochronology (Zachos et al., 2001b). Consequently, age data and chronstratigraphy of the Oligocene–Miocene boundary and the Antarctic strata that contain the boundary are also reviewed. Data sets considered to indicate climate and ice-sheet variability across the boundary include benthic and planktic oxygen isotope (d18O) records (Kennett and Shackleton, 1976; Miller et al., 1991; Wright and Miller, 1992; Paul et al., 2000; Zachos et al., 2001b; Billups et al., 2002) and microfossil geochemistry (Billups and Schrag, 2002; Lear et al., 2004), sequence stratigraphic analyses of Antarctic (Fielding et al., 1997; Naish et al., 2001, 2008) and mid-latitude (Kominz and Pekar, 2001; Pekar et al., 2002; Miller et al., 2005; Pekar and DeConto, 2006; Pekar and Christie-Blick, 2008) continental margin strata, Antarctic palaeobotany and palynology (Askin and Raine, 2000; Barrett, 2007) and physical properties of Antarctic drill core strata including lithology, clay mineralogy, mudrock geochemistry and magnetic mineralogy (Verosub et al., 2000; Ehrmann et al., 2005; Passchier and Krissek, 2008). Finally, the cause of the Mi1 glaciation is considered. 9.1.1. Identiﬁcation of the Oligocene–Miocene Boundary Early deﬁnition of the Oligocene–Miocene boundary relied on the identiﬁcation of the last occurrence of the calcareous nannofossil Dictyococcites

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bisectus (23.7 Ma; Berggren et al., 1985). However, this has proved problematic in the colder waters, coarser sediments and hiatus prone strata of the Antarctic and Southern Ocean. The reassignment of the boundary by Cande and Kent (1992, 1995) to the slightly older base of Magnetic Polarity subchron C6Cn.2n (23.8 Ma; Fig. 9.2) has made its identiﬁcation more straightforward in Antarctica and the Southern Ocean but only in relatively complete and continuous stratigraphic successions (Wilson et al., 2002; Roberts et al., 2003). More recently, the recognition of astronomically inﬂuenced cyclical physical properties and d18O records in continuously deposited deep successions has enabled astronomical calibration of late Oligocene through early Miocene time. The astronomical calibration suggested that, while still coincident with the base of subchron C6Cn.2n, the boundary was in fact nearly a million years younger (22.970.1 Ma, Shackleton et al., 2000, Pa¨like et al., 2004; 23.03 Ma, Billups et al., 2004, Gradstein et al., 2004; Fig. 9.2). The climatic signiﬁcance of this was outlined by Zachos et al. (2001b) and Pa¨like et al. (2006) who recognised the coincidence of the Oligocene–Miocene boundary and the Mi1 isotope excursion with an unusual coincidence of low eccentricity and low-amplitude variability in obliquity of the Earth’s orbit (Fig. 9.3). This would have placed the Earth in a sustained period of unusually low seasonality (cold summers), which Zachos et al. (2001b) claimed would have limited polar summer warmth and encouraged ice growth at the poles. Equally, within a few hundred thousand years, the coincidence of increased eccentricity and highamplitude variability in obliquity would have resulted in warmer polar summers and increased summer melt.

9.2. Proxy Records 9.2.1. The Isotopic Record Oxygen isotope ratios (d18O) in foraminiferal tests from deep-sea sedimentary records have long been recognised to represent the Cenozoic climatic (temperature, sea level and ice volume) history of the Earth (e.g. Shackleton et al., 1977 and references therein). However, deciphering the climatic history of Antarctica from d18O values alone in deep-sea records has always proven difﬁcult due to the ambiguity of inﬂuence on the signal from the volume of ice on land versus isotopic fractionation, which is related to the water temperature during the precipitation of calcite (Miller et al., 1991). Early studies attempted to separate the two inﬂuences by focusing their analyses on

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foraminiferal species, such as benthic forms, known to live in water masses thought to be relatively stable in their temperature history (Shackleton and Kennett, 1975; Kennett, 1977), hence deducing that any shorter-term ﬂuctuation was due to ice volume rather than temperature fractionation. While these studies recognised major threshold changes in the climatic deterioration of Antarctica, assuming no Northern Hemisphere ice sheets at the time, their resolution was limited and the record incomplete across the Oligocene–Miocene boundary. In a higher-resolution compilation of Oligocene–Miocene d18O records, Miller et al. (1991) recognised a relatively transient ca. 2 myr period and 1 m amplitude cyclicity in Miocene isotope values (isotope events Mi1-Mi7). The 1 m shifts, they suggested were of the same order as the threshold shifts identiﬁed by earlier studies (Kennett and Shackleton, 1976) and represent similar volumes of ice accumulation on the Antarctic craton. The most signiﬁcant of these shorter order isotopic events, Mi1, was coincident with the Oligocene–Miocene boundary (Fig. 9.3). Originally deﬁned by Miller et al. (1991) from DSDP Site 522 (Figs. 9.1 and 9.3), the event has subsequently been conﬁrmed at numerous locations and the timing, duration and magnitude reﬁned (Zachos et al., 1997; Paul et al., 2000; Zachos et al., 2001a, b; Billups et al., 2002, 2004, Billups and Schrag, 2002). For at least 2 myr prior to the Oligocene–Miocene boundary, d18O values were relatively stable and of low-amplitude variability (o0.5 m; Paul et al., 2000). In contrast, the Mi1 event represents a dramatic B1 m increase in d18O, over a 250 ky period immediately prior to the Oligocene–Miocene boundary peaking coincident with the boundary (Billups et al., 2004; Fig. 9.3). Peak values persisted for only B20 ky before returning to similarly low amplitude but slightly increased late Oligocene mean d18O values over the ﬁrst B120 ky of the early Miocene (Paul et al., 2000). The covariance of the isotope signal in both benthic and planktic species in the late Oligocene at Equatorial Atlantic ODP Site 929 led Paul et al. (2000) to conclude that the variability was primarily ice volume driven. However, the Mi1 event, itself, is of relatively lower amplitude in planktic records which led Paul et al. (2000) to suggest that only 0.5 m is likely due to ice-volume effects, which, using the late Pleistocene calibration, represents growth of an ice sheet in Antarctica of similar proportion to the present-day East Antarctic Ice Sheet (EAIS). The remaining 0.5 m, they concluded was due to a 21C cooling of bottom waters at ODP Site 929 in the western Equatorial Atlantic. Another approach to deciphering temperature versus ice-volume components of the deep-sea d18O signal was employed by Lear et al. (2004) who determined palaeotemperature independently from Mg/Ca ratios in foraminifera tests across the Oligocene–Miocene boundary at ODP Site 1218 in

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the eastern Equatorial Paciﬁc (Figs. 9.1 and 9.3). The 2–31C of cooling predicted in the equatorial Atlantic by Paul et al. (2000) immediately prior to Oligocene–Miocene boundary was conﬁrmed in the Equatorial Paciﬁc by Lear et al. (2004). However, peak cooling preceded peak d18O values at ODP Site 1218 and a slight (11C) warming in bottom waters of the Equatorial Paciﬁc is recorded immediately prior to and during peak Mi1 d18O values followed by another 21C of warming post-Mi1. A similar temperature variation was observed B800 ky prior to the Oligocene–Miocene boundary. Lear et al. (2004) concluded that the Mi1 event did indeed represent a signiﬁcant increase in continental ice volume and, given the warming inﬂuence predicted by Mg/Ca ratios, perhaps a larger ice-volume growth than that predicted by Paul et al. (2000). Billups and Schrag (2002) also suggested that the d18O record from ODP Site 747 represented an ice-volume signal because paired Mg/Ca measurements suggested little change in ocean temperature through the early Miocene. Recent work, however, suggests caution when interpreting stable intervals in Mg/Ca ratios from deep-water sites due to potential saturation of carbonate, which might affect the partitioning of Magnesium into benthic foraminifera (Elderﬁeld et al., 2006; Lear et al., 2008). 9.2.2. Palaeo Sea Levels Sequence and seismic stratigraphy has provided a means of relating the geologic record of continental margins to global sea-level changes that are often related to ice-volume changes at high latitudes (Vail et al., 1977; Haq et al., 1987). Sea-level history is deduced by the recognition of unconformitybounded units (i.e. depositional sequences) deposited in response to a cycle of falling and rising sea level. However, determining the relative role of tectonic subsidence and uplift versus rising and falling global sea-level and hence ice-volume ﬂuctuations (glacioeustasy) is still unresolved, particularly in pre-Pleistocene records (e.g. Macdonald, 1991). Vail et al. (1977) and Haq et al. (1987) attempted to extract the glacioeustatic signal from the comparison of records from several continental margins (Fig. 9.4). However, the large amplitudes of sea level/ice volume, the limited resolution of the resulting record and the use of proprietary data to create the sea-level records, has spurred the scientiﬁc community to collect independent data to test the records of Vail et al. (1977) and Haq et al. (1987). The Oligocene– Miocene interval of these records has sparked particular interest. Both Vail et al. (1977) and Haq et al. (1987) predicted sea-level rises and falls of between B50 and 100 m in the late Oligocene. Vail et al. (1977), however,

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using two-dimensional ﬂexural backstripping. Foraminiferal biofacies and lithofacies were then used to constrain the depths of the Oligocene margin proﬁles obtained from backstripping. A eustatic fall of B40715 or 56 m apparent sea level (apparent sea level is eustasy plus the effects of water loading on the margin, Pekar et al., 2002) was estimated across the Oligocene–Miocene boundary (Fig. 9.4), which is similar to the glacioeustasy predictions from oxygen isotopic and trace metal geochemical data (Paul et al., 2000; Lear et al., 2004). Pekar et al. (2006) provided estimates of Antarctic ice volume and the resulting changes in global sea-level for the late Oligocene by applying d18Oto-sea-level calibrations to deep-sea d18O records from a number of ODP Sites. Their results indicate that the size of the Antarctic Ice Sheet increased from approximately 50% of the present-day EAIS during the latest Oligocene to as much as 25% larger than the present-day EAIS at the Oligocene–Miocene boundary. Ice volume returned to near late Oligocene size in the early Miocene (Pekar and DeConto, 2006).

9.3. Records from the Antarctic Margin The Oligocene–Miocene boundary interval was ﬁrst sampled in Antarctica at DSDP Site 270 in the Eastern Ross Sea (Fig. 9.1). Drilling recovered a succession of silty mudstones including glaciomarine sediment, which spans the Oligocene–Miocene boundary (Hayes and Frakes, 1975; Leckie and Webb, 1983). Although an abrupt lithological change with a potential hiatus at the Oligocene–Miocene boundary is noted by Leckie and Webb (1983), poor chronological resolution prevents unambiguous correlation with the Mi1 event and the earliest Miocene. The only exposed Oligocene–Miocene boundary strata reported from Antarctica crop out on King George Island (Fig. 9.1) and include the Destruction Bay Formation (Latest Oligocene) and Cape Melville Formation (earliest Miocene; Birkenmajer et al., 1985; Birkenmajer, 1987). In a recent summary of stratigraphy and facies of the succession, Troedson and Riding (2002) concluded that a signiﬁcant glacial advance occurred at the boundary and that chronological control was good enough to suggest a correlation with the Mi1 event. The facies indicate a signiﬁcant regional ice grounding event across Bransﬁeld Strait and beyond the South Shetland Islands (Troedson and Riding, 2002). Unfortunately, drilling on the shelf and slope south of the South Shetland Islands (ODP Leg 178; Barker and Carmerlenghi, 2002) did not yield any more deﬁnitive records of the Oligocene–Miocene boundary.

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Late Oligocene/early Miocene strata reported from the East Antarctic margin include Maud Rise (ODP Leg 113 sites 689 and 690; Barker et al., 1988a), the Weddell Sea margin (ODP Leg 113 Site 693; Barker et al., 1988a) and Kerguelen Plateau (ODP Leg 120 sites 747 and 748; Schlich and Wise, 1992; Fig. 9.1). The record at Maud Rise is relatively thin and comprises exclusively siliceous and carbonate ooze, although rare glacial drop stones are reported in strata from Site 689 (Barker et al., 1988a). At the Weddell Sea margin Oligocene–Miocene sediments are also ﬁne grained and include diatom mud, clay and ooze (Barker & Carmerlenghi, 2002). Oligocene–Miocene boundary sediments at Kerguelen Plateau are also carbonate ooze (Schlich and Wise, 1992), however, foraminifera preservation and age resolution were good enough at ODP Site 747 to yield a benthic oxygen isotope stratigraphy across the boundary at that site (Wright and Miller, 1992; Billups and Schrag, 2002; Fig. 9.5) and the amplitude of the Mi1 event was much reduced compared to equatorial values, with a d18O shift of only 0.3 m across the Oligocene–Miocene boundary. A strontium isotope stratigraphy has also been assembled using planktic foraminifera from ODP Site 747 (Oslick et al., 1994). Oslick et al. (1994) reported signiﬁcant increases in 87Sr/86Sr following the early Miocene Mi isotope events with a B1 myr lag. They suggested that this increase and similar subsequent stepwise increases in early–middle Miocene oceanic 87Sr resulted from changes in the glacial state of East Antarctica. Drilling in Prydz Bay (ODP Legs 119 and 188; Hambrey et al., 1991; Cooper and O’Brien, 2004) did not yield any Oligocene–Miocene age strata δ18O (per mil)

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and Hambrey et al. (1991) concluded that this was due to erosion beneath an expanded middle-Miocene ice sheet. 9.3.1. McMurdo Sound, South Western Ross Sea The Antarctic Oligocene–Miocene record is most complete in the Victoria Land Basin as recovered in the CIROS-1 and CRP-2A drill holes (Barrett, 1989, Cape Roberts Science Team, 1999; Figs. 9.1 and 9.6). As with Prydz Bay, much of the Oligocene record of the Victoria Land Basin is marked by signiﬁcant hiatuses (Wilson et al., 1998, 2000a, b; Florindo et al., 2005), however, the latest Oligocene–early Miocene is preserved in both records (Naish et al., 2001; Wilson et al., 2002; Roberts et al., 2003). On the basis of radiometric, biostratigraphic and magnetostratigraphic data, Wilson et al. (2002) placed the Oligocene–Miocene boundary at 183.7 m in the CRP2A core at the base of a normal polarity interval correlated with Polarity Subchron C6Cn.2n using Berggren et al.’s (1995) time scale. However, following the astronomical revision of the late Oligocene through early Miocene time scale (Billups et al., 2004; Gradstein et al, 2004; Pa¨like et al., 2004), Naish et al. (2008) placed the boundary at 130.27 m in an unconformity in the CRP-2A core and revised the age of strata underlying the unconformity to encompass Polarity Chron C7n. Roberts et al. (2003) placed the Oligocene–Miocene boundary at 247 m in the CIROS-1 core B35 km south of CRP-2A. However, following the revision of Naish et al. (2008), the Oligocene–Miocene boundary in the CIROS-1 core more likely occurs in an unconformity at 92 m (Fig. 9.7). The boundary immediately overlies an unconformity at 248.71 m, which might represent as much as 1 myr following the age revision of Antarctic shelf diatom zones (Scherer et al., 2000) implied by Naish et al. (2008). The strata recovered in both the CRP-2A and CIROS-1 drill holes are cyclic in nature and interpreted to represent periodic advance and retreat of ice across the Antarctic margin concomitant with sea-level fall and rise, respectively (Fielding et al., 1997; Naish et al., 2001). Each sequence is organised into a vertical succession, which begins with an erosion surface and is followed by a diamictite and sandstone, which gives way to sparsely fossiliferous bioturbated mudstone representing a cycle of glacial advance and retreat followed by open water conditions across the site of deposition (Naish et al., 2001) in concert with changes in relative sea-level (Dunbar et al., 2008; Fig. 9.6). Naish et al. (2008) estimated the glacioeustatic inﬂuence on relative water depth changes by deconvolving the tectonic, isostatic and palaeobathymetric components of water depth. These results are consistent with the d18O

Figure 9.6: Environmental proxy data for the upper part of the CRP-2A drill core. Grain size and clast data, and sequence stratigraphic, palaeobathymetric, depositional environment and ice margin interpretations are from Cape Roberts Science Team (1999). CIA, chemical index of alteration (data from Passchier and Krissek, 2008). ARM, anhysteretic remanent magnetization (data from Verosub et al., 2000). Temperature and meltwater indicators are discussed in Barrett (2007). Nothofagus leaf in CIROS-1 was identiﬁed at 215.5 mbsf immediately underlying the Oligocene–Miocene boundary as identiﬁed from the revised age model presented in this paper (Fig. 9.7).

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sea-level calibration of Pekar et al. (2002) from the New Jersey margin. Each cycle is interpreted to represent between 10 and 40 m of eustatic variation in the late Oligocene with perhaps 50 m of sea-level fall concomitant with an ice sheet some 20% larger than present coincident with the unconformity, which is correlated with the Oligocene–Miocene boundary by Naish et al. (2008). Grounding of ice across the site at the Oligocene–Miocene boundary is also conﬁrmed by macro- and micro-structures indicative of glacio-tectonic deformation (Passchier, 2000; van der Meer, 2000). The Oligocene–Miocene boundary is marked by some signiﬁcant changes in physical properties in the CRP-2A core. Late Oligocene sedimentary cycles underlying the boundary are 55–60 m thick and relatively complete, whereas early Miocene sedimentary cycles are 10–20 m thick and truncated (Cape Roberts Science Team, 1999; Fig. 9.6). The clay mineralogy of the strata across the Oligocene–Miocene boundary in CRP-2A records stable physical weathering conditions (Ehrmann et al., 2005). Major element ratios derived from mudrock geochemistry for the same strata show signiﬁcant shifts in the chemical index of alteration (CIA) across the Oligocene– Miocene boundary, which indicate periods of increased physical weathering and mechanical erosion associated with glacial advance (Passchier and Krissek, 2008). The CIA data reported by Passchier and Krissek (2008; Fig. 9.6) was corrected for the presence of primary volcanic detritus in order to reﬂect the palaeoclimatic record within the mudrock geochemistry. Short-lived glacial events at B23, B21 and B19 Ma indicated by the CIA data are correlated with the Mi events by Passchier and Krissek (2008) and interpreted to represent signiﬁcant climatic and ice-sheet events in East Antarctica. Magnetic properties also show a marked change across the Oligocene–Miocene boundary at 130.27 m in the CRP-2A core (Verosub et al., 2000). An earlier change in magnetic properties at 270 m (late Oligocene) is attributed to inception of the McMurdo Volcanic Province (Verosub et al., 2000). Despite these changes in physical properties across the Oligocene– Miocene boundary in the CRP-2A core, palynological data although sparse due to low concentrations of organic matter, indicate a partially open landscape dominated by small Nothofagus (Southern Beech) stands or sparse tundra vegetation persisting through the late Oligocene–early Miocene (Askin and Raine, 2000). Oligocene–Miocene strata in the CIROS-1 core also contain similar amounts of pollen (Mildenhall, 1989) and a Nothofagus leaf fossil was preserved in latest Oligocene strata of the CIROS-1 Core (Hill, 1989; Figure 9.7). Mean summer temperature records derived from the (KþNa)/Al ratios of the CRP cores (Passchier and Krissek, 2008) indicate relatively constant mean summer

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temperatures of B101C in the latest Oligocene dropping to B61C in the early Miocene. Marine palynomorphs, however, suggest a more signiﬁcant change following the Oligocene–Miocene boundary with a signiﬁcant reduction in the occurrence of prasinophyte algae, which is taken to indicate a reduction in offshore meltwater inﬂuence and hence cooler climates (Barrett, 2007).

9.4. Possible Drivers of Change Across the Oligocene–Miocene Boundary 9.4.1. Atmospheric Carbon Dioxide While there has been no speciﬁc simulation of the inﬂuence of atmospheric CO2 on ice-sheet growth at the Oligocene–Miocene boundary, DeConto and Pollard (2003a) demonstrated the potential link through simulations across the Eocene–Oligocene boundary. Their modelling demonstrated that icesheet inception occurred below a threshold of 3 pre-industrial atmospheric CO2 levels. Model results also demonstrated a strong response of ice volume to orbital forcing as atmospheric CO2 approaches the glaciation threshold, and decreasing orbital variability of an established ice sheet as CO2 approaches pre-industrial levels. Despite a predicted rewarming to preOligocene–Miocene boundary levels in the early Miocene and a major Antarctic glaciation in the middle Miocene (Zachos et al., 2001a), this was not matched by a parallel changes in levels of atmospheric CO2 as determined by geochemical proxies (Pagani et al., 1999, 2005; Pearson and Palmer, 2000; Fig. 9.8). The apparent decoupling between Miocene temperatures and atmospheric CO2 levels led Pagani et al. (1999, 2005) to conclude that, despite a signiﬁcant decrease in atmospheric CO2 from B500 ppmv in the late Oligocene to new near-modern values at the Oligocene–Miocene boundary, changing atmospheric CO2 levels may have been secondary in driving Miocene Antarctic climatic and ice-sheet evolution.

9.4.2. Ocean Circulation/Tectonic Isolation The progressive opening of oceanic gateways (Fig. 9.1) and progressive tectonic isolation during the Cenozoic stages of Gondwana breakup have been indicated as critical threshold events in the climatic deterioration and inception of ice sheets since the ﬁrst deep-sea oxygen isotopic records were

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Figure 9.8: Potential controls on climate and ice sheet behaviour across the Oligocene–Miocene boundary and origin of the Mi-1 d18O event. Obliquity and eccentricity orbital target data from Laskar et al. (2004). d18O data are from Zachos et al. (2001b) and pCO2 data are from Pagani et al. (2005). All data are plotted against the astronomical time scale presented by Billups et al. (2004). See text for discussion.

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recovered (Shackleton and Kennett, 1975; Kennett, 1977). Testing of this hypothesis has proven particularly difﬁcult due to uncertainties in the timing of gateway opening and inception of deep-water circulation. While estimates for the timing of the opening and deepening of the Tasmanian Gateway between Australia and Antarctica are reasonably well constrained to the Eocene–Oligocene boundary (Stickley et al., 2004), estimates for the opening of a deep Drake Passage range from the middle Eocene (B50 Ma) to the late Miocene (B6 Ma) (Barker and Burrell, 1977; Barker et al., 2007; Livermore et al., 2007). However, some estimates suggest that deep-water circulation through the Drake Passage may well have been coincident with the Oligocene– Miocene boundary (Barker and Burrell, 1977, 1982). Pekar et al. (2006) and Pekar and Christie-Blick (2008) suggested, however, that Southern Ocean water masses were still relatively poorly mixed in the late Oligocene–early Miocene and that individual records used to compile the composite Cenozoic oxygen isotope curve by Zachos et al. (2001a) were drawn from water masses with different temperature and salinity histories. Hence, introducing an artifact from the splice of different isotopic records (Pekar and Christie-Blick, 2008) previously interpreted to represent signiﬁcant oceanic warming in the latest Oligocene (Zachos et al., 2001a). A number of numerical ocean modelling studies (Mikolajewicz et al., 1993; Nong et al., 2000; Toggweiler and Bjornsson, 2000; Sijp and England, 2004) have shown that the opening of a deep circum-Antarctic passage can cool the Southern Ocean by 1–31C. While the amount of cooling in these studies is somewhat dependent on modelling details associated with the treatment of the atmosphere (Huber and Nof, 2006), the effects of this range of cooling on continental climate and ice-sheet mass balance have been shown to be small relative to the effects of the falling Cenozoic CO2 concentrations (DeConto and Pollard, 2003a; Huber et al., 2004). For example, recent model simulations testing the importance of sea ice feedback on Antarctic Ice Sheets show that the continental interior is relatively insensitive to changes in Southern Ocean sea surface temperatures, and the effect of even large changes in ocean heat transport and sea ice is generally limited to the continental margins (DeConto et al., 2007). Conversely, the expansion of the EAIS, as presumed to have occurred at Mi1, has a dramatic effect on simulated Southern Ocean sea surface temperatures and sea ice distributions via the ice sheet’s inﬂuence on regional temperatures and low-level winds (DeConto et al., 2007; Fig. 9.9). As these simulations clearly show, a growing Mi1 ice sheet would have cooled Southern Ocean sea surface temperatures by several degrees, pushing the 0 1C isotherm equatorwards and increasing the area, thickness, and fractional cover of seasonal and perennial sea ice (DeConto et al., 2007). Furthermore, as the katabatic wind ﬁeld increased in

a

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Figure 9.9: South polar seasonal temperatures, sea ice and winds in response to a growing ice sheet as simulated by a GCM (DeConto et al., 2007). (a–c) Climatic conditions in a pre-Mi1 world with isolated ice caps in the continental interior (Top). (d–f) Climatic conditions with a fully glaciated east antarctica as presumed to have existed at the time of Mi1 (bottom). With the exception of ice sheet geometry, boundary conditions are identical in both simulations including the same late Palaeogene palaeogeography, 2 pre-industrial CO2 (560 ppmv), and a relatively cold austral summer orbit conducive to Antarctic Ice Sheet growth. Ice sheet geometries are taken from prior GCM-ice sheet simulations of Antarctic glaciation (DeConto and Pollard, 2003a). Austral Summer (December, January, February) and winter (June, July, August) seasonal climatologies are shown on the top and bottom of (a–f), respectively. (a, d) Seasonal surface (2 m) air temperature, (b, e) seasonal sea ice extent and thickness in metres and (c, f) lowest level (sigma level ¼ 0.189) GCM winds with vector scale length equivalent to 21C per m1 of wind velocity.

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intensity, the enhanced polar easterlies and westerlies would have increased ocean frontal divergence and upwelling, with possible implications for the marine carbon cycle and CO2 drawdown (DeConto et al., 2007). Such mechanisms have been implicated as important contributors to the dynamics of Quaternary glacial cycles (Stephens and Keeling, 2000; Archer et al., 2003), but they have yet to be considered in a Miocene context. While discussions of tectonic inﬂuences on Antarctic climate evolution usually focus on Southern Ocean gateways, Miocene ice sheets could have also been sensitive to changes in tropical climate associated with Tethyan tectonics. As the Southern Ocean gateways were widening and deepening, the eastern Tethys was closing. Ocean modelling studies have shown that the progressive closure of the Tethys affected the location of deep-water formation and the thermohaline component of the meridional overturning circulation, ocean heat transport, and both tropical and high-latitude sea surface temperatures (Hotinski and Toggweiler, 2003; von der Heydt and Dijkstra, 2006). While the Antarctic interior appears to have been relatively insensitive to changes in the Southern Ocean, the modern Antarctic interior receives much of its moisture from the low mid-latitudes and signiﬁcant changes in the tropics and associated teleconnections to polar latitudes could be important. Considering the timing of Tethyan closure relative to Antarctic Ice Sheet expansion in the Miocene, the sensitivity of ice-sheet evolution to low-latitude versus circum-Antarctic sea surface temperatures should be tested in future modelling studies. The perspective provided by numerical climate modelling suggests falling greenhouse gas concentrations around the time of the Oligocene–Miocene boundary (Pagani et al., 2005) had a greater impact on Antarctic climate than the direct, physical effects of ocean gateways. However, the indirect effects of the gateways, including their inﬂuence on the marine carbon cycle and atmospheric CO2 should also be considered. These indirect effects may be found to be more important to Cenozoic climate events like Mi1 than the direct inﬂuence of the gateways on ocean circulation and heat transport (Mikolajewicz et al., 1993; DeConto and Pollard, 2003a; Huber et al., 2004). 9.4.3. Orbital Parameters If varying orbital parameters were to result in ice growth on the Antarctic craton, this was most likely to have occurred when summer insolation is minimised either by the seasonal timing of aphelion (precession) during periods of relatively high eccentricity, periods of low eccentricity which minimised the

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effects of precession, even when perihelion occurs during austral summer, or during periods of reduced amplitude in obliquity. All of these orbital conﬁgurations result in relatively cool summers and hence reduce the potential for summer melt. Zachos et al. (2001b) measured oxygen isotope ratios from Ocean Drilling Programme Equatorial Atlantic (Ceara Rise) Site 929 for a B5 myr interval spanning the Oligocene–Miocene boundary. The Mi1 event was found to correlate with both a minima in the low frequency (400 ky) eccentricity cycle and a prolonged minima in the amplitude of obliquity, a co-occurrence with a reoccurrence interval of 2.4 myr or longer (Pa¨like et al., 2004, 2006). This sustained period of unusually low seasonality (cold summers) was, however, relatively transient with warmer summers returning within a few hundred thousand years from the coincidence of increased eccentricity and high-amplitude variability in obliquity which would have resulted in warmer polar summers and increased summer melt (Zachos et al., 2001b). 9.4.4. Ice-Sheet Hysteresis Coupled climate–ice sheet models have been reasonably successful in simulating sudden Cenozoic glaciation events such as Oi1 and Mi1 (DeConto and Pollard, 2003a, b). For example, beginning with an ice-free continent and assuming gradually declining greenhouse gas concentrations and accounting for orbital forcing, DeConto and Pollard (2003a) simulated the sudden stepwise glaciation of East Antarctica within a 200-ky interval. The simulated ice sheet was comparable in volume to the modern EAIS, but signiﬁcantly smaller than the volume of Mi1 ice reconstructed from the proxy isotope and sea-level records discussed above. Subsequent modelling work, including a representation of ice shelves not included in their earlier simulations (Pollard and DeConto, 2007), have shown that an Antarctic Ice Sheet B20–25% bigger than today would have required a glaciated West Antarctica, and ice grounding lines extending close to the continental shelf break around much of the margin. An Mi1 ice sheet of this size would have been similar in geometry to the ice sheet that existed at the Last Glacial Maximum (Huybrechts, 2002). However, the presumably warmer ocean at Mi1 might have been unconducive to the seaward migration of grounding lines, so this scenario maybe difﬁcult to reconcile from a modelling perspective. Furthermore, the cold south polar conditions implied by such an Ice Sheet also implies global temperatures low enough to have allowed signiﬁcant glaciation in the Northern Hemisphere, especially during orbital periods which produced cold boreal summers. While Greenland may have

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contained some glacial ice as early as the Eocene (Eldrett et al., 2007), the Oligocene–Miocene boundary is B20 myr before the onset of the ﬁst signiﬁcant Northern Hemisphere glacial cycles. Clearly, some important problems remain in terms of reconciling the magnitude of the Mi1 event. While the rapid growth of Antarctic ice at Mi1 can be explained through a combination of decreasing greenhouse gas concentrations and orbital forcing (with other possible inﬂuences from mountain uplift and/or ocean circulation), the ephemeral nature of the event and subsequent variability of ice volume are also problematic from a modelling perspective (Pollard and DeConto, 2005). As shown in both simple and sophisticated numerical icesheet models, (Weertman, 1961; Huybrechts, 1994; Pollard and DeConto, 2005), the high albedo and elevation of large polar-centred ice sheets produce considerable hysteresis. In a scenario of cooling climate, a polar ice sheet can grow suddenly, once the snow line intersects sufﬁcient land area in mountains and high plateau. The non-linear jump in ice volume is facilitated by height–mass balance and albedo feedbacks, as the ice sheet spreads horizontally (albedo feedback) and more of the parabolic ice surface rises above the snow line and out of the ablation zones around its margins (height–mass balance feedback) (Abe-Ouchi and Blatter, 1993; DeConto and Pollard, 2003a). The high elevation and albedo of the ice sheet inhibit the ice sheet from disappearing during subsequent warming interval, unless temperatures (snow lines) rise far above their initial values (elevation) at the time of glacial onset (Huybrechts, 1994). Pollard and DeConto (2005) studied this hysteresis effect in a coupled GCM–ice sheet model and in a simple ﬂowline model with parameterised mass-balance forcing. They concluded that the hysteresis effect is strong enough to preclude orbital forcing from driving the range of Cenozoic ice-volume variability seen in the oxygen isotope and sea-level records described above, unless the orbital forcing is accompanied by signiﬁcant changes in greenhouse concentrations. During favourable (cold austral summer) orbital periods, the atmospheric CO2-glaciation threshold for Antarctica is B2 pre-industrial levels, while CO2 must approach B4 pre-industrial levels during a warm austral summer orbital period to trigger the collapse of the interior EAIS. If the sensitivity of the models to orbital and greenhouse gas forcing is reasonable, the short duration of the peak Mi1 event would require a signiﬁcant perturbation to the carbon cycle, producing signiﬁcant global warming soon after the peak glacial interval. Greenhouse gas variability of this magnitude is not evident in existing proxy reconstructions of early Miocene CO2 (e.g. Pagani et al., 2005), however, higher-resolution records will be required to resolve this type of CO2 variability across key climatic events like the Oligocene–Miocene boundary.

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9.5. Summary and Conclusions Strata recovered from the Antarctic margin indicate a signiﬁcant glacial advance at the Oligocene–Miocene boundary reaching the south Shetland Islands on the Antarctic Peninsula (Troedson and Riding, 2002) and grounding in Prydz Bay and the South Western Ross Sea as indicated by hiatuses in drill cores (Hambrey et al., 1991; Roberts et al., 2003; Naish et al., 2008). Ice rafted as far north as Maud Rise (Barker et al., 1988a, b) and the central Ross Sea (Leckie and Webb, 1983) but did not appear to reach the Kerguelen Plateau (Schlich and Wise, 1992). Pre-Oligocene–Miocene boundary strata indicate a late Oligocene Antarctic Ice Sheet (Cape Roberts Science Team, 1999), which expanded to an ice volume of the order of 20% greater than the present ice sheet at the Oligocene–Miocene boundary (Naish et al., 2008). The glacial expansion, however, although signiﬁcant in extent and volume, must have been relatively transient and neither cold nor extensive enough to extinguish Nothofagus tundra vegetation (Askin and Raine, 2000; Roberts et al., 2003), which persisted across the boundary despite a slight drop in temperature (Passchier and Krissek, 2008). Marine palynomorphs, however, indicate that coastal temperatures did not return to the warmth of the late Oligocene with a much reduced freshwater melt input to coastal regions (Barrett, 2007). Data from the Antarctic continent are entirely consistent with the shortlived (200 ky) ice-volume increase from 40% of present Antarctic ice volume to 25% greater than present Antarctic ice volume across the Oligocene– Miocene boundary with concomitant oceanic deep-water cooling implied by the Mi1 isotopic excursion recognised in equatorial and Southern Hemisphere deep-sea sedimentary records (Paul et al., 2000). Ice-volume estimates are conﬁrmed by the backstripped stratigraphic records from the New Jersey Margin (Kominz and Pekar, 2001; Pekar et al., 2002), however, accommodating this much ice on Antarctica when global temperatures were presumably warmer than today may prove difﬁcult from a modelling perspective. Warm summer mean temperatures were re-established soon after the Oligocene–Miocene boundary, although a few degrees cooler than pre-Miocene summer mean temperatures. The duration and transience of the Mi1 glacial expansion and swift recovery in Antarctica likely resulted from the limited polar summer warmth from coincidence of low eccentricity and low-amplitude variability in obliquity of the Earth’s orbit at the Oligocene– Miocene boundary (Zachos et al., 2001b). This was followed by warmer polar summers and increased melt from increased eccentricity and highamplitude variability in obliquity in the early Miocene, allowing the recovery of vegetation on the Antarctic craton. Atmospheric CO2 concentrations

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remained below the 2 pre-industrial threshold, which promoted sensitivity of the climate system to orbital forcing during cold Austral summers. While climate and ice-sheet modelling supports the fundamental role of greenhouse gas forcing punctuated by orbital forcing as a likely cause of events like Mi1 (DeConto and Pollard, 2003a; Huber et al., 2004; DeConto et al., 2007; Pa¨like et al., 2006), the models underestimate the range of orbitally paced ice-sheet variability recognised in early Miocene isotope and sea-level records unless accompanied by signiﬁcant ﬂuctuations in greenhouse gas concentrations (Pollard and DeConto, 2005). While, tectonic inﬂuence may have been secondary, they may well have contributed to oceanic cooling recorded at the Cape Roberts Project site in the South Western Ross Sea (Barrett, 2007).

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Developments in Earth & Environmental Sciences, 8 F. Florindo and M. Siegert (Editors) r 2009 Elsevier B.V. All rights reserved DOI 10.1016/S1571-9197(08)00010-4

Chapter 10

Middle Miocene to Pliocene History of Antarctica and the Southern Ocean Alan M. Haywood1,, John L. Smellie2, Allan C. Ashworth3, David J. Cantrill4, Fabio Florindo5, Michael J. Hambrey6, Daniel Hill2, Claus-Dieter Hillenbrand2, Stephen J. Hunter2,1, Robert D. Larter2, Caroline H. Lear7, Sandra Passchier8 and Roderick van de Wal9 1

School of Earth & Environment, University of Leeds, Leeds LS2 9JT, UK Geological Sciences Division, British Antarctic Survey, High Cross, Madingley Road, Cambridge CB3 0ET, UK 3 Department of Geosciences, North Dakota State University, Fargo, ND 58105-5517, USA 4 Royal Botanic Gardens Melbourne, Private Bag 2000, Birdwood Avenue, South Yarra, Victoria 3141, Australia 5 Istituto Nazionale di Geoﬁsica e Vulcanologia, Via di Vigna Murata 605, 00143 Rome, Italy 6 Institute of Geography & Earth Sciences, University of Wales, Aberystwyth, Ceredigion SY23 3DB, UK 7 School of Earth, Ocean and Planetary Sciences, Cardiff University, Main Building, Park Place, Cardiff CF10 3YE, UK 8 Department of Earth and Environmental Studies, Mallory Hall 252, Montclair State University, Montclair, NJ 07043, USA 9 Institute for Marine and Atmospheric Research Utrecht, Utrecht University, Princetonplein 5, 3584 Utrecht, The Netherlands 2

ABSTRACT This chapter explores the Middle Miocene to Pliocene terrestrial and marine records of Antarctica and the Southern Ocean. The structure of the chapter makes a clear distinction between terrestrial and marine records as well as proximal (on or around Antarctica) and more distal records (Southern Ocean). Particular geographical regions are identiﬁed that reﬂect the areas for which the majority of palaeoenvironmental and palaeoclimatic information exist. Speciﬁcally, Corresponding author. Tel.: þ44(0)113 343 8657; Fax: þ44(0)113 343 6716;

E-mail: [email protected] (A.M. Haywood).

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the chapter addresses the terrestrial sedimentary and fjordal environments of the Transantarctic Mountains and Lambert Glacier region, the terrestrial fossil record of Antarctic climate, terrestrial environments of West Antarctica, and the marine records of the East Antarctic Ice Sheet (EAIS), the West Antarctic Ice Sheet (WAIS) and the Antarctic Peninsula Ice Sheet (APIS), as well as the marine record of the Southern Ocean. Previous and current studies focusing on modelling Middle Miocene to Pliocene climate, environments and ice sheets are discussed.

10.1. Introduction The Middle-to-Late Miocene interval is believed to represent a time of signiﬁcant ice-sheet expansion on Antarctica (e.g. Miller et al., 1991, 2005; Lear et al., 2000; Turco et al., 2001; Billups and Schrag, 2002; Shevenell et al., 2004). The stable isotope record of the deep sea demonstrates that a midMiocene ‘‘climatic optimum’’, at B15 Ma, was followed by strong enrichment in oceanic d18O and climatic cooling over the next 6 Ma (e.g. Zachos et al., 2001). During this interval, the East Antarctic Ice Sheet (EAIS) is thought to have been a major and permanent ice sheet, although ﬂuctuations in the size of EAIS may still have occurred (e.g. Westerhold et al., 2005). Denton et al. (1984) proposed that during this time, the EAIS overrode the Transantarctic Mountains (Fig. 10.1). Recent studies from the western Dry Valleys indicate that the atmosphere cooled by as much as B201C prior to an erosional event that is linked to the EAIS overtopping the mountains (Lewis et al., 2007). If such a sequence is repeated elsewhere (i.e. cooling and development of cold-based alpine glaciers preceded ice-sheet overriding), then it would suggest that the initial rise in deep-sea benthic d18O (Zachos et al., 2001) reﬂects deep-water cooling, followed later by ice-sheet expansion. Furthermore, Lewis et al. (2006) propose that immense freshwater ﬂoods to the Southern Ocean from large subglacial lakes beneath the expanded EAIS occurred between 14.4 and 12.4 Ma. The discharges are not only considered to have been the erosive force that formed such prominent Dry Valleys features as the Labyrinth, but are also considered through their impact on oceanic circulation to be a cause of mid-Miocene climatic change. A different explanation for the enrichment in oceanic d18O is that it represents the onset of signiﬁcant glaciation on West Antarctica (Mercer, 1978; Ciesielski et al., 1982). This is supported by the ﬁrst occurrence of ice rafted debris (IRD) at Deep Sea Drilling Project (DSDP) Site 325 in the Bellingshausen Sea during the Early to Middle Miocene

Figure 10.1: Map showing geographical locations discussed in the text. Antarctic bed topography also shown from the BEDMAP dataset (Lythe et al., 2000).

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(Hollister et al., 1976). However, the relative role that the EAIS and West Antarctic Ice Sheet (WAIS) played in the Middle Miocene Climate Transition (MMCT) has recently been questioned by new seismicstratigraphic data from the Ross Sea revealing at least ﬁve major intervals of grounded ice advance and retreat in the Middle Miocene (Bart, 2003; Chow and Bart, 2003). Much of this ice was sourced from West Antarctica, suggesting the presence of a large and dynamic WAIS prior to the MMCT (Bart, 2003; Chow and Bart, 2003). The stability of Antarctic Ice Sheets during the Late Miocene and Pliocene has been the subject of almost continuous debate for more than 20 years. The key questions in this argument are when did the EAIS switch from being polythermal and dynamic to predominantly cold based and stable, and could relatively short-lived climatic warm intervals have been sufﬁcient to inﬂuence the Antarctic Ice Sheet? A long-standing view is that the EAIS became stable during the Middle Miocene, evidence for which is primarily derived from apparent landscape stability and well-dated surfaces and ash deposits in the Dry Valleys region along the western border of the Ross Sea (e.g. Denton et al., 1993; Marchant et al., 1993a; Sugden, 1996). An alternate view is that terrestrial glacial deposits, known as the Sirius Group, scattered throughout the Transantarctic Mountains, indicate dynamic ice-sheet conditions even during the Pliocene. This conclusion is based on the occurrence of Pliocene (and older) diatoms reworked into glacial deposits (Harwood, 1983, 1986; Webb et al., 1984, 1996; Harwood and Webb, 1986; Wilson, 1995; Wilson et al., 1998). The dynamic nature of the Pliocene Antarctic Ice Sheet is supported by the Pagodroma Group along the ﬂanks of the Lambert Glacier (e.g. Hambrey and McKelvey, 2000a,b). Each view is internally consistent and scientists have been presented with a signiﬁcant challenge in reconciling the different views. When considering the size and diverse landscapes of Antarctica, we should not be surprised to see a degree of heterogeneity in the climate and environmental response. Yet the current state of knowledge is so contradictory that the scientiﬁc community has become polarized into two camps (‘‘stabilists and dynamicists’’) over the issue of Middle Miocene to Pliocene conditions on Antarctica (e.g. Harwood and Webb, 1998; Stroeven et al., 1998a,b).

10.2. East Antarctic Terrestrial Environments Two main pre-Quaternary sedimentary sequences recording glacial events are found on land in East Antarctica, the Sirius Group in the Transantarctic

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Mountains and the Pagodroma Group in the Prince Charles Mountains ﬂanking the Lambert Glacier (Fig. 10.2). The two sequences are quite different in character. The Sirius Group is a mainly terrestrial glacigenic succession whose age has proved highly controversial, with Pliocene, midMiocene or even older ages suggested. The Pagodroma Group is an iceproximal fjordal sequence up to several hundred metres thick, and is well dated on the basis of in situ marine fossils. All these well-exposed deposits are important in unravelling the history of the EAIS, as they are the nearest

Figure 10.2: (a) Map of Antarctica showing key sites of investigations of on-shore pre-quaternary glacigenic sediments. (b) The distribution of known Sirius Group outcrops. (c) Location of the Pagodroma Group sites along the Lambert Graben (redrawn from Hambrey and McKelvey, 2000a; Hambrey et al., 2003).

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deposits to the heart of the ice sheet, and yield valuable information concerning palaeogeography and palaeoclimate. 10.2.1. Transantarctic Mountains Outcrops of glacigenic sedimentary rocks of the Sirius Group occur as erosional remnants scattered throughout the Transantarctic Mountains as far south as 861S (Fig. 10.3). The deposits are exposed in two typical settings: (1) as thin erosional remnants at high elevation in palaeovalleys or on ﬂat mountain summits, or (2) as sequences more than 100 m thick along the walls of broad trunk valleys occupied by large outlet glaciers draining the modern EAIS. The main sedimentary facies are massive diamicts, massive boulder gravels, stratiﬁed pebbly sands and muds, and laminated sand/mud couplets with dropstones (Fig. 10.3). Diamicts overlie spectacular grooved and striated pavements in some areas, notably at Roberts Massif. In the Dominion Range, bordering Beardmore Glacier, well-preserved Nothofagus leaves, stems and roots, plus mosses and insects are interbedded with the diamict (Francis and Hill, 1996; Ashworth et al., 1997). The diamicts are dominated by subangular to subrounded striated and facetted clasts. The environment of deposition is interpreted as largely subglacial to proglacial, with extensive wet-based, probably polythermal glaciers being the agent of sediment transport, although a marine inﬂuence is evident through marine fossil remains in part of the Beardmore Glacier succession (e.g. Mercer, 1972; Mayewski and Goldthwait, 1985; McKelvey et al., 1991; Webb et al., 1996; Stroeven and Prentice, 1997; Wilson et al., 1998, 2002; Hambrey et al., 2003). The lithostratigraphy and depositional setting of the Sirius Group at two of the southernmost locations is summarized in Table 10.1. A debate was ignited when Webb et al. (1984), based on recycled marine diatoms, linked the Sirius Group to glacial expansion after early Pliocene deglaciation and marine incursion of East Antarctica. In contrast, Denton et al. (1984) attributed the Sirius Group to mid-Miocene overriding of the Transantarctic Mountains based on geomorphological evidence from the Dry Valleys. A dynamic ice-sheet hypothesis and a stable ice-sheet hypothesis developed, which represented contrasting views of Neogene Antarctic climate and glacial dynamics. Pivotal to this debate is our understanding of the uplift and erosional, hence unrooﬁng, history of the Transantarctic Mountains (e.g. Webb et al., 1984; Behrendt and Cooper, 1991; Sugden et al., 1995; Kerr and Huybrechts, 1999). The Transantarctic Mountains are divided into several crustal blocks which vary in size on a range of scales and which probably experienced uplift at different rates and

Figure 10.3: Glacial erosional features and sedimentary characteristics of the Sirius Group, Transantarctic Mountains. (A) Aerial view of sub-Sirius Group glacially abraded surface on Roberts Massif, showing patch of diamict, postdepositional faults, and a superimposed recessional quaternary moraine system. (B) Multiple beds of diamict comprising the W100 m thick Shackleton Glacier formation at the type locality alongside upper Shackleton Glacier. (C) Large-Scale Glacial Grooves on Jurassic Dolerite, predating deposition of the Sirius Group. (D) Striated dolerite surface with overlying Sirius Group diamict, Roberts Massif. (E) Typical massive diamict facies of the Sirius Group, showing alignment of clasts, interpreted as basal till, Roberts Massif. (F) Laminated silt and thin diamicts with large dropstone, interpreted as an ice-contact lake deposit in the Shackleton Glacier formation, Bennett Platform (redrawn from Hambrey et al., 2003).

Battye Glacier (mid-Miocene)

Bardin Bluffs (Pliocene–?Pleistocene)

Mount Johnston (?Oligocene–early Miocene)

Fisher Bench (mid-Miocene)

Fisher Massif

Formations

Amery Oasis

Massive and stratiﬁed diamictite; sandy breccia/conglomerate; partially siliciﬁed wood fragments; well lithiﬁed

Massive and weakly stratiﬁed diamict, massive boulder gravel, stratiﬁed sandstone, breccia; sand/ mud laminate with dropstones at Shackleton Glacier. Well preserved Nothofagus ﬂora at Beardmore Glacier. All facies well indurated

Principal facies

Massive boulder gravel; massive diamict; gravel; laminate with dropstones. Rich diatom microﬂora. All weakly to well indurated

Principal facies

Shackleton erosion surface/dominion erosion surface

Cloudmaker (Pliocene–?Miocene)

Shackleton Glacier

Eroded remnants as erratics

Meyer Desert (Pliocene)

Beardmore Glacier

Bennett Platform

Shackleton Glacier

Formations

Ice-proximal fjordal, grounding-line fan complexes predominant, to distal glaciomarine

Paleoenvironment

Subglacial with extensive mass-movement and ﬂuvial reworking; welldeveloped ﬂora

Subglacial; some supraglacial; proglacial glacioﬂuvial; ice contact lake

Paleoenvironment

Ages of formations are given where determined using diatom biostratigraphy. No correlation is implied in the Sirius Group. Data from Webb et al. (1996) for the Sirius Group at Beardmore Glacier; Hambrey et al. (2003) for the Sirius Group at Shackleton Glacier; Hambrey and McKelvey (2000a) for the Pagodroma Group.

Pago-droma

Group

Sirius

Group

Table 10.1: Representative stratigraphy of the Sirius and Pagodroma Groups with principal lithofacies and interpretation of palaeoenvironment. 408 A. M. Haywood et al.

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during different periods, principally in the late Cretaceous and early Tertiary (prior to c. 40 Ma; Stump and Fitzgerald, 1992; Fitzgerald, 1994; Bussetti et al., 1999; van der Wateren et al., 1999). The uplift histories of adjacent blocks might also be different. The younger uplift and erosional history is only relatively well known for the Dry Valleys block, as a result of detailed detrital studies by the Cape Roberts Project (CRP). Those studies showed that at the deepest levels cored, equivalent to c. 34 Ma, the kilometre-thick Kirkpatrick Basalts (Jurassic) had already been almost completely removed and the sand modes were dominantly derived from sandstones of the Victoria Group, then the Taylor Group (Devonian–Triassic Beacon Supergroup; Smellie, 2000a,b, 2001a,b, unpublished; Talarico et al., 2000; Sandroni and Talarico, 2001). By c. 33 Ma, the 2 km-thick Beacon Supergroup had been cut through to expose outcrops of basement rock (early Palaeozoic and Precambrian granitoids and metamorphic rocks), which then began to contribute signiﬁcant detritus. Between that time and c. 29 Ma, tectonic stability and, presumably, little uplift-related erosion are inferred from the essentially unchanging detrital modes. Further changes in the detrital modes between 24 and 16 Ma suggest another phase of instability and possibly uplift/unrooﬁng, after which the record is not preserved. Although the detrital record is interpreted here in terms of simple uplift and unrooﬁng, an alternative explanation is that the changes observed might reﬂect varied phases of climate-related erosion independent of uplift (e.g. enhanced downcutting by glaciers; cf. Kerr and Huybrechts, 1999). This is an ambiguity that has yet to be resolved. According to the dynamic ice-sheet hypothesis, the diatom assemblages incorporated in the Sirius Group record periods when East Antarctic basins were ice-free and became inundated by the sea (Harwood, 1986; Harwood and Webb, 1998). The time intervals of which diatoms are lacking record the stages when either ice was covering the inland basins or the ﬂoors of the basins were exposed. According to stabilists, however, the basic assumption derived from ice-sheet modelling is that deglaciation requires considerable climatic warming (Huybrechts, 1993). A substantial body of internally consistent evidence (including ash and cosmogenic dating) for a pre-middle Miocene landscape, subsequently unmodiﬁed by ice, has been published in numerous papers (e.g. Marchant et al., 1993a,b,c, 1996; Sugden et al., 1995; Sugden, 1996; Sugden and Denton, 2004). For example, isotopic ages of W14 Ma for the in situ, unweathered, ash deposits from the higher elevated regions of the Dry Valleys (Marchant et al., 1993b) are in disagreement with the concept of warming causing major deglaciation in the early–midPliocene. The diatom evidence for a Pliocene age of the Sirius Group has been disputed (Barrett, 1996; Gersonde et al., 1997; Stroeven et al., 1998a,b).

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Field studies and studies of aerial photographs provided evidence of rejuvenation of faulting during and after deposition of the Sirius Group (Hambrey et al., 2003). Moreover, based on morphostratigraphic constraints and provenance analyses, several authors pointed out that the Sirius Group comprises deposits of multiple glaciations within different wet-based icesheet drainage systems, which operated during consecutive stages of glacial denudation in the Transantarctic Mountains (e.g. Mercer, 1968, 1972; Brady and McKelvey, 1979, 1983; McKelvey et al., 1991; Stroeven, 1997; Van der Wateren et al., 1999; Passchier, 2001, 2004; Hambrey et al., 2003). Unfortunately, the absolute ages of the individual formations within the Sirius Group remain inconclusive and lithostratigraphic correlations are hampered by the complex tectonic framework of the Transantarctic Mountains. Given the considerable linear geographical spread of the Sirius Group over a distance of c. 1,500 km and spanning nearly 101 of latitude, the grouping of these strata under a single group name has exacerbated the arguments concerning the age. Resolution of the age question remains a key challenge for Antarctic geologists, but it is likely that, although the Sirius Group may contain disputed Pliocene or Miocene elements, it could also extend back to Oligocene time in view of the presence of glaciomarine sediments of this age offshore (Barrett, 1996; Hambrey et al., 2002; Francis et al., 2008). 10.2.2. Lambert Glacier Region The Pagodroma Group is a succession of massive diamicts and boulder gravels with minor stratiﬁed diamicts, laminites, sand and gravel, cropping out in the Prince Charles Mountains along the western margin of the Lambert Graben (Figs. 10.2 and 10.4). The best known outcrops are at Amery Oasis and Fisher Massif (Bardin, 1982; Hambrey and McKelvey, 2000a; McKelvey et al., 2001; Whitehead et al., 2003, 2006a), but equivalent exposures are known from Mount Menzies, the last exposed rock in this region heading polewards. The total outcrop distance of known glacigenic strata along the ﬂank of the graben is about 800 km. The graben is occupied by the Lambert Glacier, a south–north-ﬂowing outlet glacier, whose total ice-drainage area represents 13 per cent or 1 million km, of the EAIS. It has acted as a conduit for ice ﬂowing into Prydz Bay since its initial formation at the Eocene–Oligocene transition (Barron et al., 1991; Strand et al., 2003). Although the Pagodroma Group is regarded as the East Antarctic equivalent of the Sirius Group, there are major differences in lithofacies, facies thickness and geometry, fossil content and palaeoenvironment. For example, these

Figure 10.4: Glacial erosional features and sedimentary characteristics of the Pagodroma Group, Prince Charles Mountains. (A) Uplifted segment of a palaeofjord bottom and sidewall (arrowed), ﬁlled with nearly 300 m of iceproximal fjordal sediment of the Miocene Fisher Bench formation, Fisher Massif. (B) Striated pavement on Amery Group Sandstone, Radok Lake, Amery Oasis; this surface is coeval with the lower part of the Pliocene Bardin Bluffs formation. (C) Typical massive to crudely stratiﬁed boulder gravel of the Miocene Fisher Bench formation, Fisher Massif, interpreted as icecontact deposits associated with a grounding-line fan. (D) Weakly stratiﬁed diamict and boulder gravel of the Pliocene Bardin Bluffs formation at the type locality, interpreted as ice-proximal deposits; the cliff face here is about 60 m high. (E) Laminated silt and thin diamictites representing deposition adjacent to a grounding-line fan as quiet water cyclopels and cyclopsams, interrupted by gravity-ﬂow deposition, Miocene Battye Glacier formation, Dragons Teeth, Amery Oasis (redrawn from Hambrey and McKelvey, 2000a).

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deposits contain in situ diatom assemblages that can be used to date the formations directly (Whitehead et al., 2003, 2004). The four formally deﬁned formations within the Pagodroma Group (Table 10.1) have age ranges from early Miocene (or possibly Oligocene) to Pliocene–Pleistocene (Hambrey and McKelvey, 2000b; Whitehead et al., 2003, 2006a). The strata at Mount Menzies (Whitehead and McKelvey, 2002) have not yet yielded any datable fossils. The massive diamicts and boulder gravels indicate deposition in an ice-proximal environment near a grounding line, whereas the stratiﬁed facies represent more distal iceberg deposition. The diamicts contain striated and facetted clasts within a matrix that is ﬁner grained than modern Antarctic tills. Dropstone structures occur in thin laminated beds and resemble glacigenic marine rhythmites (cyclopels and cyclopsams) that commonly occur to the sides and front of a groundingline fan (Fig. 10.5). The depositional environments are considered to be analogous to the modern fjords of East Greenland with fast-ﬂowing polythermal tidewater glaciers (Hambrey and McKelvey, 2000b). The Battye Glacier Formation contains diatom-bearing beds (opal contents up to 15 wt.%) with in situ mollusc assemblages, which are assumed to have been deposited in ice-distal settings (Whitehead et al., 2006a). The Pagodroma Group has a cumulative thickness of W800 m and provides evidence for major shifts in the position of the grounding line of the Lambert Glacier throughout the Neogene (Hambrey and McKelvey, 2000a). The Pagodroma Group represents remnants of a much more extensive fjord-ﬁll sequence that has been uplifted above sea level (a.s.l.) (Fig. 10.4). Palaeofjord ﬂoors have been uplifted to different levels at Amery Oasis and Fisher Massif, the highest being nearly 1,500 m on the latter. In contrast, the main Lambert trough has been excavated to 2,000 m below sea level, the sediment from which has been delivered into Prydz Bay since the Eocene/ Oligocene transition, where it forms the prograded continental shelf, including the huge late Neogene Prydz Trough-Mouth Fan that was drilled during Ocean Drilling Program (ODP) Leg 188 (Passchier et al., 2003; O’Brien et al., 2004). The Prydz Bay continental margin has been drilled by ODP Legs 119 and 188 (see Barron et al., 1991; Cooper et al., 2004; Whitehead et al., 2006b). Numerical modelling of ice-sheet expansion and recession in the Lambert Glacier–Prydz Bay region has demonstrated that when major phases of erosion and sediment delivery took place, the ice dynamics were strongly controlled by the changing bathymetry of the iceeroded graben (Taylor et al., 2004). The clay mineralogy of the Pagodroma Group indicates a dominance of illite and chlorite in the Middle Miocene or older Mount Johnston and Fisher Bench Formations, indicative of physical weathering and erosion of

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Figure 10.5: Distribution of Neogene (15–2 Ma) terrestrial volcanic outcrops in West Antarctica. local metavolcanic and gneissic source rock terrains under glacial conditions (Ehrmann et al., 2003). In contrast, the Middle–Late Miocene Battye Glacier Formation contains signiﬁcant contributions of smectite and kaolinite (Whitehead et al., 2006a), and the Pliocene–Pleistocene Bardin Bluffs Formation has the highest kaolinite concentrations in the absence of smectite. Formation of kaolinite and smectite in source rock terrains generally requires prolonged exposure of source rocks to chemical

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weathering. The chemically altered nature of the strata is conﬁrmed by bulk chemical and rock magnetic studies (Bloemendal et al., 2003; Passchier and Whitehead, 2006). Recycling of weathered materials from local sedimentary source rocks cannot be ruled out (Ehrmann et al., 2003). The origin and signiﬁcance of the chemically weathered materials in the Middle Miocene and younger Pagodroma Group formations remains inconclusive at this time and further ﬁeld and laboratory studies are necessary to resolve this issue. The two regions described above represent the best studied examples of pre-Quaternary onshore glacigenic successions, but there are others elsewhere, e.g. the Antarctic Peninsula (Hambrey et al., 2008) and the Grove Mountains (Fang et al., 2004). When a better-dated Sirius Group can be linked to the offshore record in McMurdo Sound, and the data from the Prince Charles Mountains fully integrated with the emerging Prydz Bay record, we will have a good understanding of the behaviour of the EAIS since its inception in these segments of the Antarctic margin. 10.2.3. The Terrestrial Fossil Record of East Antarctic Climate The onset of Antarctic glaciation during the Oligocene dramatically altered terrestrial environments ultimately resulting in widespread extinction of plant and animal taxa (Cantrill, 2001). Tundra environments, based on pollen from the Cape Roberts borehole, are known to have existed in the Ross Sea region through the Oligocene until at least the early Miocene (Askin and Raine, 2000; Raine and Askin, 2001). The ﬁnal extinction of plants and animals with more complex life histories is assumed to be associated with the shift to the polar desert climate which prevails today. Whether or not this immediately followed the mid-Miocene warm interval or was as late as the mid-Pliocene is the most outstanding problem in Neogene climatic history in need of resolution. Pre-Pleistocene glacigenic sequences (Sirius and Pagodroma Groups) are patchily exposed throughout the Transantarctic and Prince Charles mountains. They are dominated by glaciomarine fjordal sediments, lodgement tills and outwash deposits (see previous section for more details). At the Oliver Bluffs on the upper Beardmore Glacier, fossiliferous sequences occur within the Meyer Desert Formation. The discovery of wood (Carlquist, 1987) and later Nothofagus leaves (Hill and Truswell, 1993; Webb and Harwood, 1993; Francis and Hill, 1996; Hill et al., 1996) from a location about 500 km from the South Pole makes this arguably the most important fossil site in Antarctica. The fossils provide an unusual window

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415

into Neogene environments and biogeography but until questions surrounding their age can be better resolved, their value as palaeoclimatic proxies is lessened. The deposits have been assigned a Pliocene age using a biostratigraphic argument based on reworked marine diatoms (Harwood, 1986; Webb et al., 1996). Recently, however, cosmogenic dates obtained on boulders from moraines that post-date the Meyer Desert Formation indicate ages of greater than 5 Ma (Ackert and Kurz, 2004). The authors propose various age models based on different rates of weathering and discuss different amounts of uplift, but based on their discussion, a Miocene age for the Meyer Desert Formation is preferred over a Pliocene age (Ackert and Kurz, 2004). If the older age for the Meyer Desert Formation prevails, then it is tempting to speculate that the degree of melting of the ice sheet that resulted in fjords opening up into the Transantarctic Mountains was associated with the midMiocene climatic optimum. That interpretation would be more acceptable to researchers in the Dry Valleys who present evidence that the shift from wetto cold-based glaciation occurred during the Middle Miocene (Denton et al., 1993; Marchant et al., 1993a; Sugden, 1996). The fossil Nothofagus leaves from the Meyer Desert Formation were described as a new species, N. beardmorensis, but with morphological similarities to extant species in South America and Tasmania (Hill et al., 1996). Originally, a krummholz growth form (scrubby, stunted growth form) was assigned to the Nothofagus (Webb and Harwood, 1993) but later studies of the growth ring anatomy (Francis and Hill, 1996) suggested a low prostrate and spreading habitus, more similar to ground-hugging shrubs growing today along exposed parts of treeline on Isla Navarino at the southern tip of South America (Ashworth and Kuschell, 2003, p. 193). The prostrate growth forms implied colder growing conditions than estimated using the krummholz interpretation, and considerably lower than the comparison to the cold-temperate rain forest of southern South America made on the basis of a pollen study (Askin and Markgraf, 1986; Mercer, 1986). Nearest living relative comparisons, along with physiological experiments into cold tolerance of Nothofagus, yielded theoretical estimates of mean summer month temperatures of at least 51C for 3 months and low winter month temperatures of 15 to 221C. Mean annual temperature (MAT) was estimated to have been in the range of 8 to 121C (Francis and Hill, 1996; Ashworth and Kuschel, 2003) compared to an interpolated sealevel MAT at latitude 851S today of 261C (Ashworth and Kuschel, 2003). Although much of the research to date has focused on Nothofagus, the vegetation was more complex consisting of cushion plants, grasses, ranunculids (buttercups) and mosses (Ashworth and Cantrill, 2004).

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It should be emphasized that to date, only one site has been reported from the Beardmore Glacier and leaves and wood of Nothofagus have not been reported from other Sirius Group sequences although they have been discovered recently in pre-Pleistocene glacial sequences of the western Dry Valleys (Ashworth et al., unpublished). The Meyer Desert Formation also contains fossils of beetles, including two weevil species, a higher ﬂy, single species each of a freshwater gastropod and bivalve, an ostracod and a ﬁsh (Ashworth and Kuschel, 2003; Ashworth and Preece, 2003; Ashworth and Thompson, 2003; Ashworth et al., 1997). Using overlapping autecological requirements for Nothofagus, listroderine weevils and freshwater molluscs, mean summer temperatures are estimated to have been 4–51C for at least two summer months and the MAT was estimated to be 81C, similar to the temperatures estimated from the physiological requirements of the Nothofagus fossils (Ashworth and Kuschel, 2003; Ashworth and Preece, 2003). The palaeosols, which are stratigraphically higher than the fossiliferous beds, indicate polar conditions (Retallack et al., 2001), and their development may signal the end of warmer and wetter climates in Antarctica. Regardless of age, there is a question regarding the likelihood of a tundra biome re-establishing itself in the interior of Antarctica during a warm phase if it had been extirpated during an earlier cold phase. The question is important because it sets limits on how cold the climate of Antarctica could have been before the Meyer Desert Formation biota colonized the interior. Did the species disperse from refugia around the margins of Antarctica, in which case the climate had to have been continually warmer around the margins of the continent through at least the Early Miocene? Or, if they became extinct prior to the mid-Miocene climatic optimum or mid-Pliocene warm interval (B3 Ma), could they have reinvaded the continent? For Nothofagus, seed germination is considered to be critical to answering this question. Studies suggest that Nothofagus seeds cannot survive prolonged periods of immersion in salt water (Hill et al., 1996) and consequently cannot be easily dispersed across oceanic barriers. The implications of this are that components of the vegetation survived on Antarctica even during earlier glacial phases. There is a scenario, however, based on phylogenetic and palynological studies, that allows for the possibility that Nothofagus may have dispersed from Australia to New Zealand after the Tasman Sea formed (McGlone et al., 1996). Even if transTasman dispersal occurred, it would have been from land areas with similar climatic conditions that would favour successful colonization. For organisms to disperse from South America or New Zealand to Antarctica would require not only much greater distances of dispersal but for organisms to re-establish

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themselves in complex communities in a different climatic regime. This makes the long-distance dispersal scenario for recolonization of Antarctica even less probable. We note that since deglaciation at the end of the Pleistocene, climatic conditions on many of the sub-Antarctic islands have been favourable for colonization by Nothofagus, listroderine weevils and the types of freshwater molluscs found in the Meyer Desert Formation, but it has not occurred. This strongly implies that refugia were present on the margins of Antarctica, from which plants, insects and freshwater molluscs could migrate to the continental interior during a time or times of signiﬁcant climatic warming (cf. Convey et al., in press).

10.3. West Antarctic Terrestrial Environments West Antarctica contains numerous outcrops that preserve, or potentially preserve, information on Miocene and Pliocene terrestrial palaeoenvironments. All are volcanic, although several contain associated, usually minor, sedimentary strata. Because of the long time gaps between eruptions in any volcanic centre, the environmental record they preserve is typically of a coarse resolution, and they have often been ignored as a potentially important source of palaeoenvironmental information. The outcrops are often remote or inaccessible, and many have been rarely visited. Summary information on all those outcrops of Middle Miocene and Pliocene age is given in Table 10.2 (see also Fig. 10.5). The eruptive environment is relatively well understood for a few but remains uncertain for many.

10.3.1. Antarctic Peninsula Most of the outcrops in the Antarctic Peninsula have been relatively well investigated in recent years. They form two main types: outcrops dominated by a major stratovolcano (James Ross Island Volcanic Group) and outcrops formed by multiple monogenetic volcanic ﬁelds and small isolated centres (Bellingshausen Sea Volcanic Group; Smellie, 1999). James Ross Island, situated at the northern end of the Antarctic Peninsula, comprises a single large stratovolcano 1.5 km high and about 60 km in diameter, together with numerous small satellite centres. About 50 individual eruptive episodes have been identiﬁed, with most retaining information on environmental conditions between 6.25 Ma and today. All but a few show features diagnostic of glacial conditions (Smellie, 2006). The associated ice sheets had a provenance

Ames Range

Marie Byrd Land Mount Flint

Hudson Mountains

Jones Mountains

14.7–10.0; 10.3–10.0; 7.6–5.7

14.1

10–7? (imprecise) 8.5–3.7

Undated

o6.5

James Ross Island region

Ellsworth Land Snow Nunataks

6

7.5–5.4; 2.5– 2.7

Age (Ma)

Merrick Mountains

Antarctic Peninsula Alexander Island

Outcrop

K–Ar

K–Ar

K–Ar

K–Ar

40/39Ar

K–Ar

K–Ar

Dating method

Unknown

Unknown

Subaerial, subaqueous and/or glacial

Glacial (2?)

Glacial (1)

Three overlapping stratovolcanoes

Stratovolcano

Multiple small centres

Multiple monogenetic centres

Tiny degraded outcrop remnants Dissected stratovolcano and multiple satellite centres

Subaerialc; glacial? Glacial (1) and glaciomarine (interglacial)

Multiple monogenetic centres

Comments

Glacial (1 and 2)b; subaerial

Environmental characteristicsa

LeMasurier and Thomson (1990) LeMasurier and Thomson (1990)

LeMasurier and Thomson (1990)

LeMasurier and Thomson (1990)

Nelson (1975), Skilling (1994)

Smellie and Hole (1997), Smellie (1999) Smellie (1999)

Principal sources

Table 10.2: Characteristics of volcanic outcrops of Neogene Age (15–2 Ma) in West Antarctica.

418 A. M. Haywood et al.

10.1–9.1

9.56–3.19

9.39–8.22

Toney Mountain

Hobbs Coast

Mount Murphy

5.7–4.2

10.4–8.54; 6.4

K–Ar

8.5–6.4; 7.86– 7.57

40/39Ar

K–Ar

K–Ar

K–Ar

40/39Ar

K–Ar

K–Ar

11.7–8.6

10.4–10.0; 3.0

K–Ar

13.7

Mount Cumming, Executive Committee Range Mount Sidley, Executive Committee Range Mount Bursey

Whitney Peak, Executive Committee Range Mount Hampton, Executive Committee Range Mount Hartigan, Executive Committee Range Two poorly exposed volcanoes: Lavris Peak and Tusing Peak Undissected stratovolcano Three coalesced stratovolcanoes

Subaerial?c

Subaerialc and snow contact

Glacial (1 and (mainly) 2) and subaerialc

Glacial

Unknown

Unknown

Two overlapping undissected stratovolcanoes Plateau lava basal succession Numerous small dissected centres Dissected stratovolcano

Undissected stratovolcano

Subaerial?c

Subaerial?c

Undissected stratovolcano

Subaerial?c

LeMasurier and Thomson (1990) LeMasurier and Thomson (1990) Smellie (2000a,b), LeMasurier (2002), Wilch and McIntosh (2002)

LeMasurier and Thomson (1990)

LeMasurier and Thomson (1990)

LeMasurier and Thomson (1990)

LeMasurier and Thomson (1989)

LeMasurier and Thomson (1988) Middle Miocene to Pliocene History of Antarctica and the Southern Ocean 419

2.58–2.32

4.5–3.4

Mount Berlin

Fosdick Mountains

K–Ar

K–Ar

K–Ar

40/39Ar

40/39Ar

Dating method

Three volcanoes: Mt Rees, Mt Frakes and Mt Steere Two overlapping undissected stratovolcanoes Two coalesced undissected stratovolcanoes

Glacial (2); subaerialc

Unknown

Subaerialc and glacial?

Unknown

Small monogenetic centres

Comments

Glacial (1)

Environmental characteristicsa

LeMasurier and Thomson (1990)

LeMasurier and Thomson (1990)

LeMasurier and Thomson (1990)

Smellie (2001a,b), Wilch and McIntosh (2002) Wilch and McIntosh (2002)

Principal sources

b

Ignores late-stage subaerial parasitic cones (Marie Byrd Land), often very much younger (by m.y.) than the main stratocone successions, which they overlie. Glacial 1: ‘‘thick ice’’ conditions (few hundred metres); Glacial 2:‘‘thin ice’’ conditions (o150 m); after Smellie (2000a,b). c Environmentally undiagnostic since subaerial volcanics may be supraglacial; evidence for (lower elevation) glacial conditions not exposed.

a

9.13–6.82; 8.46–5.74; 4.25–1.81 4.9–4.7

Crary Mountains

Mount Moulton

6.8–5.95

Age (Ma)

Mt. Murphy satellite centres

Outcrop

Table 10.2: (Continued ).

420 A. M. Haywood et al.

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rooted both in the Antarctic Peninsula and in a local ice cap centred on the volcano itself (Hambrey and Smellie, 2006). The ice was wet based and erosive throughout, although there are few thermal regime data for units younger than 2 Ma. Conversely, a few of the volcanic units showed evidence for eruptions in marine conditions, and marine fossils in interbedded sedimentary deposits also suggest that warmer ice-poor conditions (interglacials) occurred during at least three broad periods: 6.5–5.9, 5.03–4.22 and o0.88 Ma (Smellie et al., 2006a,b). The James Ross Island Volcanic Group is now known to contain numerous sedimentary interbeds, mainly varieties of diamicts with a glacial association that are only beginning to be investigated fully (Hambrey and Smellie, 2006; Hambrey et al., 2008). By contrast, the Bellingshausen Sea Volcanic Group contains multiple small volcanic centres with ages ranging between 7.5 and 2.5 Ma (Smellie, 1999). The glacial environment was wet-based, temperate or subpolar, and varied between thin and thick ice conditions (o70 and W500 m, respectively), although not all outcrops yielded the full range of information (Table 10.2; Smellie et al., 1993; Smellie and Hole, 1997). There also exists a small outcrop of 2.5 Ma tephra (Hornpipe Heights Formation, northern Alexander Island) erupted under essentially dry subaerial conditions, presumably when any associated glacial cover was below the present outcrop (i.e. below 750 m a.s.l.). Scant evidence from elsewhere on Alexander Island suggests that glacial conditions may have been present at 2.5 Ma, although few Pliocene samples are dated and the age resolution is poor (Smellie, 1999). 10.3.2. Ellsworth Land Three outcrops occur in this region: Snow Nunataks, Jones Mountains and Hudson Mountains (Table 10.2). The Hudson Mountains are particularly poorly known. Conversely, Snow Nunataks comprises several small undated nunataks that have been divided into the Mount Benkert and Mount McCann formations, formed of subaqueous stratiﬁed hyalotuff and subaerial lava, respectively (Smellie, 1999). The formations are cogenetic and resemble basaltic glaciovolcanic centres known as tuyas (Smellie, 2000a,b) erupted probably under ice thicknesses that locally exceeded 350 m. However, a marine environment, while probably less likely, cannot be excluded since so little is known about the area. A 700 m-thick Cenozoic volcanic succession resting on a subhorizontal polished and striated glacial unconformity crops out at Jones Mountains (Rutford et al., 1972; Hole et al., 1994). Diamictite lenses interpreted as tillite overlie the unconformity and the

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overlying sequence has been divided into at least two units of pillow lava, hyaloclastite breccia and hyalotuff. Striated and facetted erratic boulders are also scattered throughout the lower volcanic unit. The descriptions most closely resemble volcanic sequences formed under a thin glacial cover. However, since individual ‘‘thin ice sequences’’ are theoretically limited to thicknesses r150 m (Smellie, 2000a,b, 2001a,b), the much greater total thickness of volcanic rock present suggests that more than two volcanic sequences are present. This is also suggested by the presence of numerous reworked hyalotuff lenses dispersed within the sequence, and possibly by the wide range in the published isotopic ages (mainly 7–10 Ma; Table 10.2), which might suggest an outcrop constructed from multiple eruptions that occurred over a long period. 10.3.3. Marie Byrd Land Marie Byrd Land contains the largest Cenozoic volcanic ﬁeld in Antarctica (LeMasurier and Thomson, 1990). It is also the most remote region and is among the least visited and least well known of all Antarctic volcanoes. Numerous reports published in the 1970s and 1980s suggested a subglacial origin for many centres (summarized by LeMasurier and Rocchi, 2005). While agreeing generally that the eruptive environment was probably glacial, the published criteria used most often (e.g. mainly presence of hyaloclastite; less often interbedded tillite and/or glacially scoured surfaces) are too simplistic and permit only limited interpretation of the eruptive environment, except in rare outcrops where more modern detailed studies have been carried out. In addition, most Marie Byrd Land volcanoes oc. 12 Ma in age have undergone only minimal dissection, probably as a result of a change to a polar thermal regime (LeMasurier and Rocchi, 2005; cf. Armienti and Baroni, 1999), so only the latest and topographically highest erupted products are available for examination. As those units were usually erupted subaerially, they provide little of environmental use except a limiting elevation for any coeval ice sheet. However, caldera and sector collapses have led to a few inland volcanoes being well exposed (e.g. Crary Mountains, Mount Sidley: Wilch and McIntosh, 2002), and many of the coastal volcanoes are deeply dissected (e.g. Mount Murphy: LeMasurier et al., 1994; Smellie, 2000a,b, 2001a,b; LeMasurier, 2002; Wilch and McIntosh, 2002; LeMasurier and Rocchi, 2005). Only the better exposed and described examples are considered here. Mount Murphy is situated on the coast in eastern Marie Byrd Land. It is a complex polygenetic ediﬁce constructed from at least three centres, of which

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only the westernmost is well exposed. The basal section, up to c. 1,400 m a.s.l., is basaltic and it has a thinner capping sequence of mainly trachyte lavas. The basaltic section is dominated by multiple units individually several tens of metres thick and composed of stratiﬁed hyalotuff, hyaloclastite breccia and sheet or pillow lava. Each unit is separated by a glacial unconformity that is often draped by a thin diamict (tillite), and a few units are capped by strombolian scoria. Although LeMasurier (2002) interpreted the Mount Murphy succession in terms of lava-fed deltas (and therefore, by implication, eruptions beneath ‘‘thick ice’’), the sequences are much more similar to those associated with a thin glacial cover, i.e. composed mainly of snow, ﬁrm and/or fractured ice that was never more than tens of metres thick (Smellie, 2000a,b, unpublished). The sequences extend down to below 400 m a.s.l., and they reﬂect the local dynamics of a small ice cap probably similar in appearance to that centred on Mount Murphy today. Conversely, several satellite centres at Icefall Nunatak, Turtle Rock and Hedin Nunatak have younger ages of 6.8–5.95 Ma (Table 10.2). They were associated with a low-elevation (o600 m a.s.l.) ice sheet composed of relatively thick ice (at least several hundred metres). That ice was thicker than today and had a surface elevation that ﬂuctuated through 200 m during the period (Smellie, 2000a,b, 2001a,b). A wet-based thermal regime (temperate or subpolar ice) characterized the glacial cover associated with the main (8–9 Ma) Mount Murphy ediﬁce, but its nature is unknown for the younger satellite centres. Finally, Mount Murphy is overlain by a thin laminated mudstone interpreted as a glacial lake sequence related to a signiﬁcant but undated overriding event (o3.5 Ma) by an extraordinarily thick former ice sheet that was at least 1,500 m thicker than at present (LeMasurier et al., 1994). Crary Mountains (9.13–1.81 Ma; Table 10.2) are situated 250 km inland of Mount Murphy, where the surrounding WAIS surface elevation is 1,600 m a.s.l. Only the two older northern centres (Mount Rees and Mount Steere) are well exposed and have ages in the range between 8 and 9 Ma, i.e. coeval with Mount Murphy. The maﬁc to intermediate successions are individually o100 m thick and composed of alternating ‘‘dry’’ lavadominated and ‘‘wet’’ hyaloclastite-rich lithofacies, interpreted as lavas interacting with local slope ice and snow (Wilch and McIntosh, 2002). Thus, the coeval Crary Mountains and Mount Murphy successions are broadly similar, being products of interaction with a thin glacial cover. They differ, however, in that the former lack the interbedded tillites and glacial surfaces that are so prominent at Mount Murphy, a difference that was tentatively interpreted to be due to the Crary Mountains ‘‘ice’’ being cold based.

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Finally, sector collapse at Mount Sidley has exposed a magniﬁcent 1,300 m-high caldera wall section that extends from c. 2,700 to W4,000 m a.s.l. The outer slopes are essentially undissected and expose only subaerially erupted volcanic products, including a widespread ‘‘dry’’ trachyte– phonolite fall deposit. By contrast, the caldera wall section is mainly formed of interbedded phonolite and trachyte lavas, domes and ignimbrite ranging in age from 5.69 to 4.18 Ma. The sequence is essentially subaerial. Evidence for glacial conditions is absent, apart from minor signs of water interaction (stratiﬁed hyalotuffs and thin vitroclastic lava bases) in one of the subsequences dated as 5.36–4.81 Ma. The latter was interpreted as due to limited glaciovolcanic interaction, presumably local slope ice/snow. However, the basal 1.5 km of the volcano is obscured by the present-day WAIS and the earlier (and lower elevation) environmental history is therefore unseen.

10.4. The Marine Record of the East Antarctic Ice Sheet Due to the nearly complete ice coverage of the Antarctic continent, a record of past Antarctic climate lies in the layers of sediments eroded over million years from the continent and deposited in sedimentary basins around its margin and in the Southern Ocean. During the past ca. 20 years, nine ODP legs have signiﬁcantly advanced our understanding of the Cenozoic tectonics and palaeoenvironments of the Antarctic region. These legs include Leg 113 in the Weddell Sea (Barker et al., 1988, 1990), Leg 114 in the sub-Antarctic South Atlantic (Ciesielski et al., 1988, 1991), Leg 119 in Prydz Bay and on Kerguelen Plateau (Barron et al., 1989, 1991), Leg 120 on Kerguelen Plateau (Schlich et al., 1989; Wise et al., 1992), Leg 177 in the southeast Atlantic sector of the Southern Ocean (Gersonde et al., 1999, 2003), Leg 178 on the Antarctic Peninsula (Barker et al., 2002), Leg 181 (Carter et al., 2000), Leg 188 in Prydz Bay (O’Brien et al., 2001; Cooper et al., 2004) and Leg 189 in the Tasmanian region (Exon et al., 2001, 2004). The pioneer of Antarctic drilling was the DSDP in 1973 (Hayes et al., 1975). Following that, four distinct phases of drilling have taken place in the Ross Sea region, in sedimentary basins under the direct inﬂuence of East Antarctica, by rigs mounted on the fast-ice that rings the southern part of the Ross Sea just north of the Ross Ice Shelf (see Chapter 3 for further details). These include the Dry Valley Drilling Project (DVDP) in 1970–1975 (McGinnis, 1981), the McMurdo Sound Sediment and Tectonic Studies (MSSTS) in 1979 (Barrett, 1986), the Cenozoic Investigations in the Western

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Ross Sea (CIROS) in 1984 and 1986 (Barrett, 1989; Barrett et al., 1992) and the multinational CRP in 1997–1999 (Cape Roberts Science Team, 1998, 1999, 2000; Hambrey et al., 1998; Barrett et al., 2000, 2001). These efforts are continuing through ANDRILL (Antarctic geological drilling) programme, which demonstrated ability to recover high quality marine and glacimarine sedimentary drill cores from high latitude ice-covered areas (Naish et al., 2007; Florindo et al., 2008; Harwood et al., 2008). To date, in spite of these efforts, important questions and problems remain unresolved and are the subject of a considerable debate within the scientiﬁc community. Among these are the causes of the mid-Miocene cooling event at around 14 Ma, and the warming–cooling climatic phases of the Pliocene. 10.4.1. Middle Miocene Cooling Marine deep-sea oxygen isotope records indicate that a period of ice-sheet instability during the Early and Middle Miocene is superimposed on a longterm warming trend that culminated in the mid-Miocene climatic optimum from B17 to 15 Ma (Miller et al., 1987; Zachos et al., 2001). This phase was followed by a gradual cooling with a major cooling event occurring at the Middle–Late Miocene boundary B14 Ma that continued for B1–2 m.y. (MMCT; Flower and Kennett, 1995). The Prydz Bay area was visited during the Leg 119 (Barron et al., 1989, 1991) and Leg 188 (O’Brien et al., 2001; Cooper et al., 2004) drilling campaigns of the ODP. Middle Miocene–Pliocene sediments were drilled at Sites 739, 742, 1165 and 1166. An interesting stratigraphic record across this transition was recovered during ODP Leg 188 when three sites were drilled proximal to the EAIS across the Prydz Bay continental shelf (Site 1166), slope (Site 1167) and rise (Site 1165). Of these, Site 1165 (64122uS; 67113uE) records a history of sedimentation on the continental rise extending back to earliest Miocene times (about 22 Ma). Several long-term changes characterize this record, including an overall trend of decreasing sedimentation rates from the bottom to the top of the hole. There is a progressive decrease in the sedimentation rate above about 308 m b.s.f. (metres below sea ﬂoor), which is marked by a transition from predominantly dark-grey ﬁssile claystones with abundant silt laminae to grey diatom-bearing clays with a higher abundance of IRD (Fig. 10.6). At this transition, the total clay content also increases. The chronology of this sequence indicates a Middle Miocene age (ca. 14.3 Ma) for the lithological transition. Correlation to ODP Hole 747A on the Kerguelen Plateau shows that this transition coincides with the base of the Mi-3/3a d18O event (Florindo et al., 2003). Middle Miocene

Figure 10.6: Lithostratigraphic column of Site 1165 and correlation between Middle Miocene intervals to ODP Hole 747A from the Kerguelen Plateau (ODP Leg 120). The lithological transition at about 308 m b.s.f., which is characterized by an increase in ice-rafted debris, sand grains and total clay content, coincides with the base of Mid-Miocene glacial event (Mi-3/3a) (modiﬁed from Florindo et al., 2003).

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changes in sedimentation at Site 1165 may have been caused by the development of a cooler ice sheet characterized by decreased rates of glacial erosion and decreased production of melt water. In such an event, less suspended sediment is delivered to the continental rise, but ice at sea level produces more icebergs and, hence, more ice rafting of debris. 10.4.2. Late Miocene–Early Pliocene Ice-Sheet Fluctuations Long-term variation in the marine oxygen isotopic data (e.g. Hodell and Venz, 1992; Kennett and Hodell, 1995; Zachos et al., 2001) are suggestive of a warming in the earliest Pliocene, culminating in the mid-Pliocene Climatic Optimum at B3 Ma. However, the oxygen isotope ratios do not allow major changes in ice volume on Antarctica. Obliquity-driven oscillations (41 kyr) in d18O of up to 0.6m, which dominate the Pliocene period, while capable of producing B20–30 m sea-level changes (depending on the temperature contribution to the d18O signal), are also considered insufﬁcient to cause large-scale deglaciation of Antarctica without a signiﬁcant increase in the temperature of the deep ocean. Marine sedimentary deposits exposed on land and on the continental shelf, slope and rise show evidence of dynamic ice-sheet behaviour in the Late Miocene–Early Pliocene and Early Pliocene warming. At Marine Plain in the Vestvold Hills, East Antarctica, a ca. 8 m thick early Pliocene sequence of diatomaceous sands and silts is exposed (Pickard et al., 1988). During deposition of the diatomaceous sediments, the ice margin must have been ca. 50 km further inland and no ﬂoating ice was covering the site. The presence of Cetacean skeletons, including dolphins and a right whale, and preliminary isotope measurements led Quilty (1993) to infer warmer conditions during deposition than currently prevail. Water temperatures may have been as high as 51C. A marked increase in sea-ice indicator diatoms and a coarsening upward of the sediments in the upper 3 m of the deposit indicate cooling conditions and glacial expansion. The sediments at ODP Sites 739, 742 and 1166 in Prydz Bay record evidence of repeated advances and retreats of the Lambert ice stream across the shelf in the Late Miocene through Early Pliocene (Hambrey et al., 1991; Passchier et al., 2003; O’Brien et al., 2004). During repeated advances of the Lambert ice stream to the shelf break glacigenic debris ﬂows build up a trough-mouth fan on the continental slope. It is worth mentioning that the IRD signal inferred from the W250-mm size fraction at Site 1165, in a sediment drift on the continental rise off Prydz Bay, suggests warmer conditions in the Early Pliocene

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(Gru¨tzner et al., 2005). The low sea-ice concentrations derived from diatom assemblages (Whitehead et al., 2005) and the warm-water periods inferred from high abundances of silicoﬂagellates (Dictyocha) at Site 1165 (Whitehead and Bohaty, 2003) are consistent with isotopic estimates of warmer conditions than present during the Early Pliocene (Hodell and Venz, 1992). In addition, three periods of enhanced accumulations of sediment originating from East Antarctica are recognized in the Late Miocene–Pliocene record of the East Kerguelen Ridge sediment drift, in the Indian Ocean to the north of Prydz Bay (Joseph et al., 2002). The sediment pulses are interpreted as periodically warmer, less stable, ice-sheet conditions during this time. DVDP-10, 11 and 15, MSSTS-1 and CIROS-2 drill-holes in the Ross Sea Region have recovered partial late Neogene records that suggest a dynamic Plio-Pleistocene history of glacial advance and retreat in the Taylor and Ferrar valleys. Of these, a 166 m thick Plio-Pleistocene sequence of diatombearing glaciomarine strata was cored in Ferrar Fiord by the CIROS-2 drilling campaign (163131uE; 77141uS) (Barrett et al., 1992; Barrett and Hambrey, 1992). The Pliocene sequence is B62 m thick and is composed of diamictite dominated by basement lithologies interbedded with thin mudstone intervals, indicating several early Pliocene advances and retreats of ice through the Transantarctic Mountains. Further evidence for Pliocene Antarctic ice volume ﬂuctuation is recorded by glaciomarine strata from the DVDP-10 and 11 drill-holes (163132uE; 77135uS) (McGinnis, 1981; Ishman and Reick, 1992; Wilson, 1993). Combined foraminiferal and magnetostratigraphic analyses indicate highly ﬂuctuating palaeoenvironmental conditions in the Late Miocene–Early Pliocene with less ice-inﬂuenced conditions in the early Pliocene (Ishman and Reick, 1992). Late Neogene trough-mouth fans have been identiﬁed in seismic data on the Ross Sea and Weddell Sea continental slopes. The Crary Fan on the Weddell Sea margin was formed by glacial sediment supply from Dronning Maud Land and ice drainage from East Antarctica through the Weddell Sea (Bart et al., 1999), whereas the Ross Sea fan was fed by ice draining through the Northern basin (Bart et al., 2000). In contrast, the geometry of Late Neogene sequences on the Wilkes Land margin is characterized as steep prograding foresets and trough-mouth fans that are less developed (Escutia et al., 2005). However, the close proximity to the deep subglacial Wilkes Land basin results in ice grounding below sea level, which may have contributed to ice-sheet behaviour that is different from the other portions of the East Antarctic margin. Interpretations of the seismic data record a dynamic Late Neogene EAIS with large expansions and contractions for the Ross Sea and Weddell Sea portions of the East Antarctic margin (Bart et al., 1999, 2000).

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10.5. The Marine Record of the West Antarctic and Antarctic Peninsula Ice Sheets The late Cenozoic glacial history of West Antarctica is archived in the sedimentary sequences of its continental margins. During the last 20 years, multichannel seismic data were acquired from the margins of the Weddell Sea, the Antarctic Peninsula, the Bellingshausen and Amundsen seas, and the eastern Ross Sea. DSDP Legs 28 (Ross Sea margin) and 35 (Bellingshausen Sea) and ODP Legs 113 (Weddell Sea) and 178 (Paciﬁc margin of the Antarctic Peninsula) recovered sedimentary sequences spanning the Neogene to Quaternary. In the following, we summarize the main results of these studies and their implications for the history of WAIS and Antarctic Peninsula Ice Sheet (APIS) from the Middle Miocene to Pliocene. Seismostratigraphic studies suggest that an early stage of glaciation characterized by presence of local ice caps and tidewater glaciers had affected the Ross Sea region already during the Late Oligocene and the Early Miocene (De Santis et al., 1995, 1999; Anderson and Shipp, 2001). Glacial and glaciomarine sediments interbedded with ﬂuvial to deep-water mudstones were deposited on the innermost part of the shelf (CIROS-1 drill site), while open-marine sedimentation dominated on the outer shelf (Abreu and Anderson, 1998; Anderson and Shipp, 2001). This depositional pattern points to cool temperate interglacial conditions, but gives no direct evidence for presence of the WAIS (Abreu and Anderson, 1998; Anderson and Shipp, 2001). Based on seismic-stratigraphic investigations and sedimentary records drilled at DSDP Sites 270 and 272, De Santis et al. (1999) inferred that a glacial transitional regime developed in the Ross Sea region during the Early and Middle Miocene. Most of the subglacial erosion and deposition still took place on the inner shelf, but occasionally local ice caps expanded onto the outer shelf. Around this time, the ﬁrst signiﬁcant amounts of IRD were deposited in the Bellingshausen Sea (DSDP Leg 35, Sites 323 and 325) and the Weddell Sea (ODP Leg 113, Sites 693 and 694) (Abreu and Anderson, 1998). The onset of IRD sedimentation at Site 325 points to an establishment of West Antarctic ice caps outside the Ross Sea region during the Early to Middle Miocene (Hollister et al., 1976; Abreu and Anderson, 1998; Barker et al., 2002). By analysing Mg/Ca ratios of calcareous benthic and planktonic foraminifera tests, Lear et al. (2000) and Shevenell et al. (2004) demonstrated that ice-sheet build-up in Antarctica between 14.2 and 13.8 Ma contributed to the pronounced, global increase in d18O ratios in benthic foraminifera tests observed by Zachos et al. (2001), even though the related eustatic sea-level fall was rather indistinct (Abreu and Anderson, 1998; Miller et al.,

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2005). While most of the Middle Miocene ice-volume increase may have occurred in East Antarctica and possibly the Arctic (Moran et al., 2006), seismic studies also point to an expansion of ice caps in West Antarctica or their coalescence to a WAIS smaller than present (Abreu and Anderson, 1998; Anderson and Shipp, 2001). In the Ross Sea region, seismic data and sedimentary records reveal that grounded ice advanced at least ﬁve times across the shelf (Chow and Bart, 2003). A comparison with foraminiferal oxygen isotope and sea-level data shows that at least two of these grounding events fall into the early part of the Middle Miocene (Bart, 2003). Bart and Anderson (1995) proposed on the basis of seismic proﬁles from the western Antarctic Peninsula outer shelf that glacial erosional surfaces formed during the Middle to Late Miocene, when huge sediments drifts started to grow on the adjacent continental rise as indicated by seismic investigations (Rebesco et al., 1997, 2002). Bart et al. (2005) combined seismic data from the western Antarctic Peninsula shelf with sedimentological data from ODP Leg 178 Site 1097 drilled on the shelf, and concluded that the earliest grounding events on the Antarctic Peninsula shelf are of Late Miocene age. Seismic studies of the sedimentary development of the Crary Fan in the southeastern Weddell Sea indicate a minimum of 5 long-term and 14 small-scale ice-sheet expansions across the adjacent shelf from the Middle Miocene to the Pleistocene (Anderson and Shipp, 2001, and references therein). However, present-day ice-drainage patterns raise the question of whether or not WAIS expansion made a major contribution to the inferred ice advances (cf. Bart et al., 1999). Despite episodic ice expansion in parts of West Antarctica during the Middle Miocene, benthic and neritic planktonic diatoms, which were reworked from the Antarctic shelf and deposited in the central Weddell Sea (Site 694), point to the absence of a major WAIS or APIS from the Middle Miocene to the early part of the Late Miocene (Kennett and Barker, 1990; cf. Chow and Bart, 2003). Widespread glacial unconformities assumed to be Late Miocene in age are observed in seismic proﬁles from the eastern Ross Sea margin (Cooper et al., 1991; De Santis et al., 1995; Anderson and Shipp, 2001). The grounded ice advances were accompanied by the deposition of glacial and glaciomarine strata (DSDP Leg 28 Site 270 and CIROS-1), which might exhibit an interior West Antarctic provenance (Anderson and Shipp, 2001) and which are interbedded with meltwater deposits and diatom oozes documenting temperate interglacial episodes (Abreu and Anderson, 1998). De Santis et al. (1999) correlated seismic data with sedimentary records drilled at DSDP Leg 28 sites and proposed an onset of fully glacial conditions in the Ross Sea region during the latest Miocene and Early Pliocene. A large and thick WAIS advanced up to the shelf edge (Accaino et al., 2005) and gave the

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Ross Sea shelf its typical modern seabed morphology, which is characterized by landward dipping, overdeepening and troughs carved by grounded ice streams (De Santis et al., 1999). Shelf progradation observed in seismic proﬁles west of the Antarctic Peninsula indicates a general expansion of APIS assumed to be Late Miocene to Early Pliocene in age (Larter and Barker, 1991; Larter and Cunningham, 1993; Bart and Anderson, 1995; Larter et al., 1997; Barker et al., 2002; Bart et al., 2005). Similar shelf progradation is reported from seismic proﬁles elsewhere on the continental margin in the Bellingshausen and Amundsen seas although the age of the prograding sequences in these areas is unconstrained (Nitsche et al., 1997, 2000; Scheuer et al., 2006). Episodic advance and retreat of grounded ice across the eastern and western Antarctic Peninsula shelf throughout the latest Miocene and Early Pliocene are evident from seismic unconformities (Sloan et al., 1995; Larter et al., 1997; Bart and Anderson, 2000; Bart, 2001; Bart et al., 2005) and the onset of shelf overdeepening west of the Antarctic Peninsula (Camerlenghi et al., 2002). Recovery of sediments deposited in subglacial and glaciomarine settings from the western Antarctic Peninsula shelf (ODP Leg 178 Sites 1097 and 1103) conﬁrm episodic APIS grounding events at least since the Latest Miocene to Early Pliocene (Barker et al., 1999a, 2002; Eyles et al., 2001). The general APIS expansion during the Late Miocene to Early Pliocene coincided with a northward shift of the Antarctic Circumpolar Current (ACC) (Herna´ndez-Molina et al., 2006), pronounced growth of large sediment drifts (Rebesco et al., 1997), clay-mineral ﬂuctuations (Hillenbrand and Ehrmann, 2005) and distinct maxima in accumulation rates of terrigenous detritus (Gru¨tzner et al., 2005) on the western Antarctic Peninsula rise. The sedimentary development of the drifts suggests episodic, enhanced supply of glacial-sourced sediments from the adjacent shelf from ca. 9.4 Ma onwards. Considerable growth of APIS and WAIS during the Late Miocene is also indicated by the onset of IRD deposition and sedimentation of a clay mineral assemblage pointing to the dominance of physical weathering in the South Orkney region (Site 696) and of rapid turbidite deposition with a West Antarctic provenance in the central Weddell Sea (Site 694) (Kennett and Barker, 1990). Anderson and Shipp (2001) infer that during the Late Miocene, the WAIS reached its modern size and, at times, was even larger than today. The seismic and core data from the western Antarctic Peninsula margin mentioned above suggest a similar scenario for the APIS. At the Miocene/Pliocene boundary, the deposition of IRD at Site 325 increased (Hollister et al., 1976), while the ﬁrst signiﬁcant deposition of IRD started at Site 322 near Drake Passage (Abreu and Anderson, 1998). Kennett and Barker (1990) attributed a shift from pelagic to hemipelagic deposition

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at Site 696 and from turbiditic to hemipelagic sedimentation at Site 694 during the earliest part of the Pliocene to the development of a stable APIS and WAIS or climatic warming in West Antarctica. Seismic stratigraphy combined with sedimentological data from DSDP Leg 28 drill cores indicates that the eastern Ross Sea shelf reached a late fully glacial stage with an overdeepened topography, which was mainly inﬂuenced by glacial processes of erosion and deposition resulting in relatively low sedimentation rates (De Santis et al., 1999). Throughout the Early Pliocene, high-frequency grounding events are observed in seismic proﬁles from the eastern Ross Sea shelf and the western Antarctic Peninsula shelf (Alonso et al., 1992; Larter and Cunningham, 1993; Larter et al., 1997; Bart and Anderson, 2000; Bart, 2001). Cyclicity in drift sedimentation recorded in drill cores from the Antarctic Peninsula rise conﬁrms that ice streams repeatedly advanced to the shelf break throughout the Pliocene (Barker et al., 2002; Hillenbrand and Ehrmann, 2005; Hepp et al., 2006). During the Early Pliocene, sedimentation rates increased in the deep Bellingshausen Sea (DSDP Leg 35 Sites 322, 323 and 325) and the Weddell Sea (ODP Leg 113 Sites, 693 and 697) (Barker, 1995). On the western Antarctic Peninsula continental rise, down-slope transport exceeded along-slope transport throughout the Pliocene (Herna´ndez-Molina et al., 2006), and the sediment supply from the shelf enhanced and maintained the sediment drifts (Rebesco et al., 1997, 2002). Seismic studies of the continental margin in the western Bellingshausen and Amundsen seas revealed truncations of foreset reﬂectors and extensive progradation, probably during the Pliocene (Nitsche et al., 1997, 2000; Scheuer et al., 2006), suggesting that grounded ice eroded the shelf and transported the terrigenous detritus across the shelf edge. According to Scheuer et al. (2006), the accumulation rate of terrigenous sediment on the adjacent continental rise reached its maximum during the Pliocene and Quaternary, which may indicate a relatively late establishment of WAIS in the adjacent hinterland (cf. Nitsche et al., 1997). In contrast, sedimentation rates on other parts of the West Antarctic margin reached distinct maxima during the Late Miocene (Site 1095; Barker et al., 2002) or during the Early Pliocene (Barker, 1995). During the Late Pliocene, sedimentation rates decreased at the drifts on the Antarctic Peninsula rise (ODP Leg 178, Sites 1095, 1096 and 1101; Barker et al., 1999a, 2002), in the deep Bellingshausen Sea (DSDP Leg 35, Sites 322, 323 and 325; Barker, 1995) and in the Weddell Sea (ODP Leg 113, Sites 694 and 697; Kennett and Barker, 1990; Barker, 1995). A concomitant transition from progradation to aggradation is observed in seismic proﬁles from elsewhere on the West Antarctic margin, from the Antarctic Peninsula to the eastern Ross Sea (Larter and Barker, 1991; Larter and Cunningham, 1993;

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De Santis et al., 1995; Larter et al., 1997; Nitsche et al., 1997, 2000). In contrast to these ﬁndings, Rebesco et al. (2006) compared changes in the geometries of seismic reﬂection proﬁles and concluded that enhanced progradation of the continental shelf and slope started almost all-around Antarctica at ca. 3 Ma. These authors suggest that APIS, WAIS and EAIS switched from a predominantly polythermal to a predominantly cold-based mode at this time of major ice-sheet build-up in the Northern Hemisphere (cf. Barker, 1995).

10.6. Marine Records of the Southern Ocean Today there are four main water masses south of the polar front: Antarctic Surface Water; Warm Deep Water (WDW); Circumpolar Deep Water (CDW); and Antarctic Bottom Water (AABW) (see Chapter 4 for further details). WDW ﬂows into the Southern Ocean from the North Atlantic, Paciﬁc and Indian Oceans. This nutrient-rich water mass upwells at the Antarctic Divergence, replacing sinking Antarctic Surface Water. CDW, the dominant water mass in the Southern Ocean, is a mixture of WDW and Antarctic waters. AABW forms as cold, saline shelf waters sink and mix with CDW prior to ﬂowing northwards out of the Southern Ocean. The Antarctic coastal current ﬂows from east to west around the continent. North of the Antarctic Divergence, prevailing westerlies drive surface waters towards the east, forming the ACC, which in places extends to the seaﬂoor. In part, these oceanographic conditions are a response to climate on the Antarctic continent. For example, formation of sea ice increases the density of shelf waters, aiding the formation of AABW. Conversely, changing ocean– atmosphere circulation patterns inﬂuence global climate by determining the transport of heat and moisture onto the polar continent. A commonly cited example is the thermal isolation of Antarctica by the ACC.

10.6.1. The MMCT The MMCT witnessed a major reorganization of global deepwater circulation patterns associated with the expansion of the Antarctic Ice Sheet (e.g. Woodruff and Savin, 1989; Wright et al., 1992; Flower and Kennett, 1995). The widespread presence of hiatuses in Southern Ocean sedimentary cores is one indicator of this Middle Miocene reorganization of ocean circulation patterns. Although these hiatuses present difﬁculties in

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reconstructing past oceanographic conditions using marine cores, a broad consensus regarding the history of some of the Southern Ocean water masses has emerged in recent years. One proxy used to reconstruct ocean circulation patterns is the carbon isotopic composition of benthic foraminifera, as water masses generally become enriched in the light isotope of carbon as they age. Interbasinal offsets in benthic foraminiferal carbon isotopes indicate that the Early Miocene Southern Ocean received warm saline deep water sourced from low latitudes in the Tethyan region (Woodruff and Savin, 1989; Wright et al., 1992). The ﬂow of this warm saline water mass likely ceased in the Middle Miocene, and the consequent reduction in meridional heat transport has been called upon as a trigger for the ice-sheet expansion in the Middle Miocene (Woodruff and Savin, 1989; Flower and Kennett, 1995). The ACC connects the Paciﬁc, Atlantic and Indian Oceans and thus plays an important role with respect to global heat transport. Inception of a deep circumpolar ACC possibly occurred near the Oligocene–Miocene boundary, well before the MMCT (e.g. Pfuhl and McCave, 2005; Lyle et al., 2007). However, a variety of different proxies all point to temporal variations in the intensity of the ACC since its inception. A reduction in surface-to-thermocline foraminiferal d18O gradients at DSDP Site 516 indicates reduced thermal stratiﬁcation in the surface waters of the southwestern South Atlantic around 16 Ma (Pagani et al., 2000). At the same time, the abundance of alkenones at this site decreased, implying a reduced nutrient input to this area. These observations have been interpreted to reﬂect a reduction in Antarctic Intermediate Water caused by a reduction in the strength of the ACC (Pagani et al., 2000). Re-establishment of a strong surface-to-thermocline temperature gradient in this region points to renewed ﬂow of the ACC coincident with the expansion of the Antarctic Ice Sheet in the Middle Miocene (Pagani et al., 2000). Because the oceanic residence time of neodymium is comparable to the ocean mixing time, the neodymium isotope system can be used as a radiogenic isotope tracer of ocean water masses. A study of the neodymium isotopic composition of sediments from DSDP Site 266 in the Australian– Antarctic basin indicates an increased contribution of neodymium from the Kerguelen volcanic province during the Middle Miocene, suggesting enhancement of the ACC during the climate transition (Vlastelic et al., 2005). Paciﬁc ferromanganese crusts that grew in equatorial Paciﬁc bottom water display a gradual change in their neodymium isotopic composition between

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38 and B20 Ma, which has been interpreted as increasing ﬂow of Southern Ocean waters into the Paciﬁc Ocean via the Deep Western Boundary Current (DWBC) as a result of progressive build-up of the ACC (Van de Flierdt et al., 2004). The isotope trends in the ferromanaganese crusts reversed in the Middle Miocene, likely reﬂecting increased mixing in the Paciﬁc as a result of increased production and export of AABW associated with the build-up of Antarctic ice (Van de Flierdt et al., 2004). Independent evidence for intensiﬁcation of the DWBC associated with the Middle Miocene Antarctic Ice Sheet growth is found in the record of sortable silt mean size from ODP Site 1123 situated on the Chatham Rise (Hall et al., 2003). The sortable silt mean size increases with stronger near bottom current ﬂow and selective deposition and winnowing. The ODP Site 1123 record indicates increased ﬂow speeds during the MMCT, interpreted as reﬂecting increased export of Southern Ocean waters to the Paciﬁc Ocean. Furthermore, the intensity of the DWBC was found to exhibit variability on the timescale of the 41 kyear orbital obliquity cycle (Hall et al., 2003). High-resolution foraminiferal proxy records from the South Tasman Rise (southwest Paciﬁc) give further independent evidence for a strengthening of the ACC associated with the ice-sheet expansion in the Middle Miocene. Shevenell et al. (2004) generated paired oxygen isotope (d18O) and Mg/Ca records using a planktonic foraminifera, G. bulloides, from ODP Site 1171. The d18O record reﬂects a combination of temperature and the oxygen isotopic composition of seawater, the latter a function of global ice volume in addition to regional salinity. The Mg/Ca record is used as a salinityindependent temperature proxy. The ODP Site 1171 Mg/Ca record displays a B71C cooling (point-to-point) between 14.2 and 13.8 Ma, which suggests that the paired d18O record reﬂects a concomitant surface water freshening above the South Tasman Rise. These records are interpreted as reﬂecting an intensiﬁcation of the ACC between 14.2 and 13.8 Ma in a series of steps paced by orbital cycles (Shevenell et al., 2004), although it has also been suggested that a component of the surface water freshening may represent massive melting pulses of the Antarctic Ice Sheet (Holbourn et al., 2005). 10.6.2. Late Miocene–Early Pliocene The Late Miocene witnessed further high latitude cooling, summarized in Hodell and Kennett (1986). A cooling of Southern Ocean upper circumpolar and intermediate water masses is documented in benthic foraminiferal oxygen isotope records from ODP Site 1088 in the sub-Antarctic South

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Atlantic, which display a two-step increase between B7.4 and 6.9 Ma (Billups, 2002). Associated with this late Neogene cooling are further changes in Southern Ocean oceanography. A 10–5 Ma hiatus in the sediment package underlying the Bounty Fan at ODP Site 1122 suggests an intensiﬁcation of the ACC, associated with the Late Miocene climate cooling and expansion of the WAIS (Carter et al., 2004). The strengthening of North Atlantic Deep Water (NADW) in the Late Miocene led to global oceanic circulation patterns approaching the modern conﬁguration (e.g. Wright et al., 1991). Late Miocene strengthening of NADW is supported by benthic foraminiferal carbon isotope records from ODP Site 1088 in the sub-Antarctic South Atlantic, which indicate an inﬂux of a nutrient-depleted water mass into the Southern Ocean at B6.6 Ma (Billups, 2002). The relative ﬂux of NADW into the Southern Ocean reached modern proportions by B6 Ma, and subsequently intensiﬁed yet further, supporting speculation that the climatic warmth of the Early Pliocene resulted from enhanced thermohaline overturn (Kwiek and Ravelo, 1999; Billups, 2002). The warm period in the Early Pliocene coincides with an apparent increase in biological productivity as indicated by increased opal deposition rates at ODP Sites 1165, 1095, 1096 and several gravity core sites in the Southern Ocean, which has been interpreted as reﬂecting a reduction in the extent of sea ice (Hillenbrand and Fu¨tterer, 2002; Gru¨tzner et al., 2005; Hillenbrand and Ehrmann, 2005; Hillenbrand and Cortese, 2006). Surprisingly, this warm interval appears to coincide with glacial advance (indicated by increased terrigenous input from Antarctica), perhaps supporting the ‘‘snowgun hypothesis’’ of Prentice and Matthews (1991) (Gru¨tzner et al., 2005). Large oscillations in the Al/Ti ratio, the lead isotopic composition and the neodymium isotopic composition of sediments at DSDP Site 266 in the Australian–Antarctic basin have been interpreted to reﬂect either variations in the intensity of the ACC or variations in weathering ﬂuxes resulting from periods of instability of the Antarctic Ice Sheet around 1.9 and 3.3 Ma (Vlastelic et al., 2005). Diatom assemblages at ODP Sites 1165 and 1166 suggest that sea-ice cover remained lower than today through much of the Pliocene (Whitehead et al., 2005), and silicoﬂagellate assemblages have been used to pinpoint three intervals (3.7, 4.3–4.4 and 4.6–4.8 Ma) within the Pliocene when sea surface temperatures in the Southern Ocean were roughly 51C warmer than today (Whitehead and Bohaty, 2003). Similar events are recorded in ODP Leg 178 Sites, pointing to circum-Antarctic warming that may have been responsible for ice-sheet loss (Escutia et al., 2007).

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10.7. Modelling Antarctic Climates and Ice Sheets Aspects of global Middle Miocene to Pliocene climates and environments have been examined in palaeoclimate modelling studies. For example, the impact that oceanic gateways such as the Panama Gateway (Klocker et al., 2005), the Indonesian Throughﬂow (Cane and Molnar, 2001) and Tethyan circumglobal passage (Hotinski and Toggweiler, 2003) had on global oceanic circulation patterns and heat/moisture transports has received considerable attention. The role of Tibetan uplift, as well as other regional orographic and palaeogeographic variations, in changing atmospheric circulation patterns and the development of monsoon systems has also been a key area of research (Prell and Kutzbach, 1992; Kutzbach et al., 1993; Fluteau et al., 1999; Bice et al., 2000; Zhisheng et al., 2001; Liu et al., 2003; Kutzbach and Behling, 2004; Ramstein et al., 2005). Reconstructing and predicting the climatic implications of altered distributions of vegetation types (Dutton and Barron, 1996, 1997; Cosgrove et al., 2002; Haywood et al., 2002a; Haywood and Valdes, 2006), sea surface temperature gradients (SST) and latitudinal temperature gradients/heat transports have been studied in detail (Chandler et al., 1994; Sloan et al., 1996; Rind, 1998; Haywood et al., 2000; Barreiro et al., 2006). However, very few palaeoclimate modelling studies have speciﬁcally addressed the question of how climates and environments changed on Antarctica during this interval. The majority of global palaeoclimate modelling experiments have used speciﬁed Greenland and Antarctic ice volumes based solely on estimates of global sea level. Therefore, modelling studies conducted thus far could be considered as representing a suite of sensitivity experiments that have tested the response of climate models to different initial conditions rather than accurate climate ‘‘retrodictions’’. One example of a palaeoclimate modelling study that did focus on Antarctica is Haywood et al. (2002b). In this study, they attempted to assess the sensitivity of a climate model to different Antarctic Ice Sheet conﬁgurations and to address, indirectly at least, the controversy existing over the behaviour of the EAIS during the Pliocene epoch and whether or not it was possible for Nothofagus remains founds in Sirius Group deposits in the Transantarctic Mountains to be Pliocene in age. Three sensitivity experiments were performed for the Pliocene using the UK Meteorological Ofﬁce’s General Circulation Model (GCM). Each experiment used a different conﬁguration of Antarctic ice conforming to a range of sea-level estimates supported by geological evidence (speciﬁcally 35, 25 and 12–15 m higher than modern sea level). Climate outputs from these experiments were used to drive

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a biome model capable of predicting the equilibrium vegetation state for a given climate forcing. For the experiments using an Antarctic Ice Sheet conﬁguration equivalent to a +12 to 15 and +25 m sea level, the biome model predicted the occurrence of an exclusively tundra type of vegetation on deglaciated regions of Antarctica (Fig. 10.7). In the +35 m experiment, climate ameliorated sufﬁciently to allow the model to predict deciduous taiga montane forests in areas of Antarctica where fossil Nothofagus material has been recovered from Sirius Group sediments (Fig. 10.7). This work suggested that it is feasible for Nothofagus to have existed in Antarctica during the Pliocene with ice-sheet extents calibrated to global sea-level records. This research had certain limitations relating to the experimental design. The conﬁguration of Antarctic ice was speciﬁed within the GCM and the resulting climatologies were only used to drive an ofﬂine biome model. A more sophisticated approach would use a coupled climate–ice and vegetation model capable of predicting, rather than being prescribed with, the equilibrium condition for Antarctic ice cover and vegetation with climate. As a step towards this, Francis et al. (2007) dynamically coupled the Topdown Representation of Interactive Flora and Foliage Including Dynamics (TRIFFID) dynamic global vegetation model to the HadAM3 GCM set up for the mid-Pliocene. The results supported the earlier study by Haywood et al. (2002b) since the TRIFFID model predicted tundra vegetation on deglaciated regions of Antarctica. A number of studies have utilized ice-sheet models to predict the response of the EAIS to increasing surface temperatures the results from which, in a few cases, have been related to the debate over how dynamic or stable the Antarctic Ice Sheet was likely to have been during the Middle Miocene to Pliocene (e.g. Barker et al., 1999b). Huybrechts (1993) used a threedimensional ice-sheet model to determine ice-sheet geometries under various kinds of climatic conditions. A surface temperature warming of between 17 and 201C was needed to generate the ice-free corridor over the Pensacola and Wilkes subglacial basins hypothesized by Harwood (1983) and Harwood and Webb (1986). For temperature rises of less than 51C, the model predicted that the EAIS increased in size due to an increase in snowfall. Even with the most favourable model set up for ice loss, a surface temperature rise of 151C was still required to make the Pensacola and Wilkes subglacial basins ice-free which Huybrechts (1993) concluded was unlikely for the interval in question. However, there are a number of limitations inherent in such an approach. First, by assessing the response of the modern EAIS to a certain climatic forcing, an assumption is made that the EAIS during the Middle Miocene and Pliocene was the same as it is today. Other important differences may

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Pliocene control experiment

Biome key:

Pliocene dyanamic experiment

Pliocene stable experiment

Figure 10.7: Predictions of Antarctic vegetation distributions for each of the Pliocene experiments using climatological means derived from the UK Meteorological Ofﬁce General Circulation Model within a biome model (BIOME 4; Kaplan, 2001).

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also have existed. For example, the Transantarctic Mountains may have been substantially lower in the past (McKelvey et al., 1991). Kerr and Huybrechts (1999) tested this possibility in an ice-sheet-modelling study and concluded that the surface elevation of individual mountain blocks has only a very local effect on EAIS dynamics. Second, by applying a uniform climate forcing within the model, the possibility for important regional-scale variations in temperature and precipitation, resulting from the differences in Middle Miocene and Pliocene boundary conditions, is ignored. Third, the ice-sheet model used did not include a representation of ice shelves or ice streams. Ice-sheet theory states that the stability of an ice sheet grounded below sea level depends on its surrounding ice shelves that are believed to buffer the ice sheet from oceanic changes (Lingle, 1984). Any weakening of these ice shelves may lead to a runaway process in which the entire marine ice sheet potentially collapses. Support for this hypothesis comes from recent observations of the terrestrial APIS. After the collapse of the major part of the Larsen A ice shelf in 1995, the ﬂow of glaciers, which had fed the former ice shelf, accelerated signiﬁcantly (De Angelis and Skvarca, 2003). The ice shelves provide a mechanism for linking ocean temperature change with the retreat of ice shelves and grounded ice-sheet instability that may have been important during the Middle Miocene to Pliocene interval. Current research is using an alternative approach to explore the nature and behaviour of the Antarctic Ice Sheet during the Middle Miocene to Pliocene interval. Instead of attempting to reconstruct the actual state of the ice sheet at a speciﬁc time, which is extremely difﬁcult given the incomplete geological record and uncertainties inherent in the techniques used for palaeoenvironmental reconstruction, modelling studies are now focusing on developing ensemble predictions which encompass the range of plausible behaviour of the ice sheet for particular periods. For the Pliocene, a suite of palaeoclimate modelling experiments which cover the range of possible boundary conditions, and which have been validated against global Pliocene palaeoenvironmental proxies, are being used to provide three-dimensional thermomechanical ice-sheet models with the required climatological forcings to predict the extent of the both the Greenland Ice Sheet and EAIS. Ranges of different initial conditions are used within ice-sheet models so that icesheet hysteresis can be explored in detail. The dependence of these results on the GCM boundary conditions and the initial state of the ice sheets leads to a range of ice-sheet reconstructions, but it seems likely that the mean state for the Pliocene Greenland Ice Sheet and EAIS were, to some extent, smaller than they are today (Fig. 10.8). For Antarctica, the model results clearly show that the Wilkes and Aurora Basins are the areas of the EAIS that are most susceptible to increased temperatures. Nevertheless, even in the most

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Figure 10.8: Examples of (A) smallest and (B) largest modelled EAIS in equilibrium with reconstructed Pliocene climate. extreme scenario, the predicted EAIS still covers all but the northernmost reaches of the Wilkes subglacial basin, which is suggested to be a source of the diatoms found within the Sirius Group sediments (Hill et al., 2007). Comparison of the model-predicted ice sheet and sea-level mean state to observed Pliocene highstands (e.g. Dowsett and Cronin, 1990) provides a potential match, which may be improved in future modelling exercises that incorporate variability in orbital forcing.

10.8. Summary 10.8.1. Terrestrial and Fjordal Sedimentary Environments of the Transantarctic Mountains and Lambert Glacier Region Two main pre-Quaternary sedimentary sequences recording glacial events are found on land in East Antarctica, the Sirius Group in the Transantarctic Mountains and the Pagodroma Group in the Prince Charles Mountains ﬂanking the Lambert Glacier. The Sirius Group is a mainly terrestrial glacigenic succession, commonly a few tens of metres thick or as a thin drape, but whose age has proved highly controversial, with Pliocene, Middle Miocene or even older ages suggested. The Pagodroma Group is an

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ice-proximal fjordal sequence up to several hundred metres thick, and is well dated on the basis of in situ marine fossils. All these well-exposed deposits are important in unravelling the history of the EAIS, as they are the nearest deposits to the heart of the ice sheet, and yield valuable information concerning palaeogeography and palaeoclimate. The two regions described above represent the best studied examples of pre-Quaternary onshore glacigenic successions, but there are others elsewhere, e.g. the Antarctic Peninsula and the Grove Mountains (Fang et al., 2004). When a better dated Sirius Group can be linked to the offshore record in McMurdo Sound, and the data from the Prince Charles Mountains fully integrated with the emerging Prydz Bay record, we will have a good understanding of the behaviour of the EAIS since its inception. 10.8.2. The Terrestrial Fossil Record of Antarctic Climate The terrestrial Neogene biota of Antarctica is poorly known with fossils coming only from exposures on the Beardmore Glacier and from boreholes in the Ross Sea region. This will soon change with the recent discoveries of Neogene fossil assemblages in the Dry Valleys (Ashworth, A. C., Lewis, A. R. and Marchant, D. M., personal communication, 2006). What is known so far is that the Neogene biota had a low diversity and was very different than the diverse forested habitats of the early Palaeogene. Evidently, tundra vegetation was present in the Transantarctic Mountains from the Oligocene until at least the Early Miocene and during that time it appears that there was a declining biodiversity (Askin and Raine, 2000; Raine and Askin, 2001). The ﬁnal extinction of several vascular plants and animals with more complex life histories is assumed to be associated with the shift to the polar desert climate which prevails today in the interior of the continent. Whether or not this immediately followed the mid-Miocene climatic optimum or was as late as the mid-Pliocene is the most outstanding problem in Neogene climatic history in need of resolution. The extinction implies that a climatic threshold was passed in which temperatures and available moisture were signiﬁcantly reduced after the event than those which supported the tundra biota. The most parsimonious interpretation to explain the occurrence of the Neogene biota is that the organisms were descendents of the Gondwana biota that had existed on the continent for millions of years, long after the fragmentation of the continent. The conclusion that plants and animals continued to inhabit Antarctica for millions of years after the continent became isolated makes their extinction

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an especially signiﬁcant event and supports the argument that it was associated with a major reorganization of the global climate system. 10.8.3. Terrestrial Environments of West Antarctica Over the past 7 Ma, the APIS has always been relatively thin (few hundred metres). It was also wet based and erosive, at least until about 2 Ma ago. There is also evidence for several ice-poor periods (interglacials) preserved in northern volcanic sequences, particularly in latest Miocene and Early Pliocene. A relatively thin ice cover (probably tens of metres) existed in Ellsworth Land, and it may have locally reached a few hundred metres, but is very poorly dated. Although there are numerous large volcanoes in Marie Byrd Land, each potentially preserving important evidence for past glacial conditions, most are essentially undissected and the only eruptive products available for study are subaerial units. They thus yield minimal useful environmental information other than maximum possible elevations for any former ice sheet(s) that may be hypothesized from other evidence. A glacial cover existed in Marie Byrd Land by 9–8 Ma, at least. It was thin (tens of metres) and formed local icecaps on the large volcanoes. It was wet based at the coast but may have been dry-based (polar) inland. A thicker ice sheet (at least several hundred metres) existed by c. 7–6 Ma, but its thermal regime is unknown. There is also evidence for at least one important episode of overriding by an unusually thick ice sheet in Pliocene or younger time (W1.5 km thick; o3.5 Ma). 10.8.4. The Marine Record of the East Antarctic Ice Sheet Marine deep-sea oxygen isotope records indicate that a period of ice-sheet instability during the Early and Middle Miocene is superimposed on a longterm warming trend that culminated in the late Middle Miocene climatic optimum from B17 to 15 Ma (Miller et al., 1987; Zachos et al., 2001). This phase was followed by a gradual cooling with a major cooling event occurring at the Middle–Late Miocene boundary B14 Ma that continued for B1–2 m.y. (Flower and Kennett, 1995). Long-term variation in the marine oxygen isotopic data (e.g. Hodell and Venz, 1992; Kennett and Hodell, 1995; Zachos et al., 2001) is suggestive of a warming in the earliest Pliocene, culminating in the mid-Pliocene Climatic Optimum at B3 Ma. However, the oxygen isotope ratios do not allow major changes in ice volume on Antarctica. Obliquity-driven oscillations (41 kyr) in

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d18O of up to 0.6m, which dominate the Pliocene period, while capable of producing B20–30 m sea-level changes (depending on the temperature contribution to the d18O signal), are also considered insufﬁcient to cause large-scale deglaciation of Antarctica without a signiﬁcant increase in the temperature of the deep ocean. 10.8.5. Marine Evidence for Antarctic Peninsula and West Antarctic Ice Sheets From marine evidence, it is possible to deduce that signiﬁcant ice-sheet growth in West Antarctica commenced during the Middle Miocene around ca. 14 Ma. In the Ross Sea and Weddell Sea regions, WAIS reached its fully glacial conﬁguration, comparable with today, during the Late Miocene. The APIS established its present size during the later part of the Late Miocene. In the western Bellingshausen and Amundsen seas, the WAIS may not have grown to its full size until the Early Pliocene. Throughout the Late Miocene and Pliocene, WAIS and APIS repeatedly advanced across the adjacent shelves, and during the Early Pliocene, WAIS and APIS may have been more dynamic (Barker, 1995; Bart, 2001; Hepp et al., 2006). Seismic proﬁles from the West Antarctic continental margin suggest there was no prolonged deglaciation during the late Neogene, but these records cannot be used to determine whether or not the WAIS and APIS decreased drastically in size during short-term interglacial periods. Investigations on sedimentary records drilled at the ODP Leg 178 sites do not corroborate a signiﬁcant ice volume reduction (Barker et al., 2002; Hillenbrand and Ehrmann, 2005). 10.8.6. Marine Records of the Southern Ocean The widespread presence of hiatuses in Southern Ocean sedimentary cores is one indicator of the reorganization of ocean circulation patterns that occurred during the MMCT, which is believed to have been associated with the expansion of the Antarctic Ice Sheet (e.g. Woodruff and Savin, 1989). Interbasinal offsets in benthic foraminiferal carbon isotopes suggest that the Early Miocene Southern Ocean received warm saline deep water sourced from low latitudes in the Tethyan region (e.g. Woodruff and Savin, 1989). The ﬂow of this warm saline water mass likely ceased in the Middle Miocene, and the consequent reduction in meridional heat transport is a possible trigger for the ice-sheet expansion in the Middle Miocene. Furthermore, a number of Southern Ocean records reconstruct an increase in the strength of

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the ACC which may also have played an important role in the MMCT (e.g. Pagani et al., 2000). During the Late Miocene, Southern Ocean records indicate further highlatitude cooling possibly associated with the establishment and/or growth of the WAIS and an intensiﬁcation of the ACC (e.g. Kennett and Barker, 1990). By 6 Ma, the relative ﬂux of NADW into the Southern Ocean reached modern proportions, and subsequently intensiﬁed yet further, supporting the speculation that the relative climatic warmth of the Early Pliocene was a result of enhanced thermohaline overturn (e.g. Kwiek and Ravelo, 1999). During the Pliocene, biological productivity increased suggesting a reduction in sea-ice extent/coverage at this time (e.g. Gru¨tzner et al., 2005; Hillenbrand and Cortese, 2006). Furthermore, silicoﬂagellate assemblages have been used to pinpoint three intervals (3.7, 4.3–4.4 and 4.6–4.8 Ma) within the Pliocene when sea surface temperatures in the Southern Ocean were roughly 51 warmer than today (Whitehead and Bohaty, 2003).

10.8.7. Modelling Antarctic Middle Miocene to Pliocene Climates and Antarctic Ice Sheets A number of palaeoclimate modelling studies have been conducted for the Middle Miocene to Pliocene period (e.g. Ramstein et al., 2005). However, only a few of them have focused on reconstructing the climate and environments of Antarctica (see Haywood et al., 2002b; Francis et al., 2007; Hill et al., 2007). A limited number of studies, using a prescribed EAIS within the climate model which is based on global sea-level estimates, suggest that during warm phases of the Pliocene (speciﬁcally the mid-Pliocene), climatic conditions of Antarctica could have ameliorated sufﬁciently to allow tundra forms of vegetation to exist in deglaciated regions. Idealized sensitivity experiments (e.g. Huybrechts, 1993), in which ice-sheet models are forced with different temperature and precipitation regimes, suggest that a surface temperature warming of between 17 and 201C is needed to generate the ice-free corridor over the Pensacola and Wilkes subglacial basins hypothesized by Harwood (1983) and Harwood and Webb (1986) for the Pliocene. For temperature rises of less than 51C, the models predict that the EAIS increased in size due to an increase in snowfall. More recent studies in which climatologies derived from globally evaluated climate modelling simulations for the Pliocene that are used ofﬂine to drive an ice-sheet model suggest that the mean state for the Pliocene Greenland Ice Sheet and EAIS were, to some extent, smaller than today. For Antarctica, the model results

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clearly show that the Wilkes and Aurora Basins are the areas of the EAIS that are most susceptible to increased temperatures. Nevertheless, even in the most extreme scenario, the predicted EAIS still covers all but the northernmost reaches of the Wilkes subglacial basin, which is suggested to be a source of the diatoms found within the Sirius Group sediments (Hill et al., 2007).

ACKNOWLEDGEMENTS This chapter is a contribution to the SCAR programme Antarctic Climate Evolution (ACE) and the BAS core science programme Greenhouse to Icehouse Evolution of the Antarctic Cryosphere and Palaeoenvironment (GEACEP). Acknowledgements for funding and logistical and technical support with the ﬁeld campaigns and drilling projects addressing Middle Miocene and Pliocene are made to UK Natural Environment Research Council, the US Ofﬁce of Polar Programs (NSF-OPP), Antarctica New Zealand, the Australian Antarctic Division, the Italian National Antarctic Program (PNRA), the Alfred Wegener Institute for Polar and Marine Research (AWI), Deutsche Forschungsgemeinschaft (DFG) and the Ocean Drilling Program.

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Developments in Earth & Environmental Sciences, 8 F. Florindo and M. Siegert (Editors) r 2009 Elsevier B.V. All rights reserved DOI 10.1016/S1571-9197(08)00011-6

Chapter 11

Late Pliocene–Pleistocene Antarctic Climate Variability at Orbital and Suborbital Scale: Ice Sheet, Ocean and Atmospheric Interactions Tim Naish1,2,, Lionel Carter1, Eric Wolff3, David Pollard4 and Ross Powell5 1

Antarctic Research Centre, Victoria University of Wellington, P.O. Box 600, Wellington, New Zealand 2 GNS Science, P.O. Box 30368, Lower Hutt, New Zealand 3 British Antarctic Survey, High Cross, Madingley Road, Cambridge CB3 OET, UK 4 Earth and Environmental Systems Institute, Pennsylvania State University, University Park, PA 16802, USA 5 Department of Geology and Environmental Geosciences, Northern Illinois University, DeKalb, IL 60115-2854, USA

ABSTRACT Continental margin drill core and seismic data indicate that between 3.0 and 2.5 Ma, high-latitude climate cooling drove both the West and East Antarctic Ice Sheets towards their present expanded cold polar state. Ice margins developed permanent marine termini with ice shelves. Direct physical sedimentary records of Antarctic Ice Sheet variability (e.g. ice-rafted debris, proximal glacimarine cycles) and more distal ocean records of sea-ice distribution (e.g. diatom palaeoecology), thermohaline circulation (e.g. sortable silt), ocean temperatures (e.g. d18O), frontal dynamics and surface circulation (e.g. palaeoecological assemblages and sea-surface temperature (SST) reconstructions) all show a strong covariance with the 41 kyr cycle in Earth’s obliquity between 3 and 1 Ma. Glacial periods are characterised by northward expansion of seasonal sea ice, SSTs up to 61C colder than now, equatorward migration of frontal zones by B5–101

Corresponding author. Tel.: þ64 4 4636197; Fax: þ64 4 4635186;

E-mail: [email protected] (T. Naish).

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latitude, intensiﬁcation of zonal westerly winds, invigorated surface circulation (e.g. Antarctic Circumpolar Current) and intensiﬁed abyssal currents. These processes lead the d18O ice volume maximum by 3–7 kyr, at which time Antarctic Ice Sheets were fully extended onto the continental shelf. Antarctic ice volume changes were in part controlled by the effect of Northern Hemisphere glacioeustasy on its marine margin, and this mechanism accounts for much of the orbital variability of the last 2.6 myr. An enigmatic interval of foraminiferal ooze and coccolith-bearing sediments in Weddell Sea and Prydz Bay sediment cores, together with a bioclastic limestone in the Ross Sea at B1 Ma, imply a signiﬁcant warming and change in ocean chemistry around the periphery of Antarctica. This event, which occurs within the short normal polarity Jaramillo Subchron in Ross Sea cores, is correlated with warm Marine Isotope Stage 31. The anomalous warming implies an increase of 4–61C in SST, possible incursion of Subantarctic Surface Waters and depression of the lysocline – an event that is apparently unique in the last 3 myr. The proximal geological record implies a second cooling step occurred at the Mid-Pleistocene Climate Transition (ca. 900 ka). Glacialinterglacial cycles, spanning the last 1 myr, recovered in drill cores from the Ross Sea and Prydz Bay are dominated by coarse-grained, ice-proximal diamicts of cold, polar subglacial afﬁnity. Open-ocean pelagites and hemipelagites are rare on the shelf. The atmospheric temperature and greenhouse gas records from Antarctic ice cores show a pronounced increase in amplitude at the 100 ka periodicity, that is coherent and in phase with marine temperature records, but lags ice volume signal. The records of greenhouse gases from the Dome C ice core show very strong congruence with many features of the temperature record, and are consistent with CO2 in particular playing a signiﬁcant role in temperature ampliﬁcation. This also suggests that the Southern Ocean plays a leading role in controlling the atmospheric concentration of CO2 on glacial-interglacial time scales. Higher dust ﬂux to Antarctica during glacial maxima compared with warm periods is considered to have radiative affects over Antarctica by providing nutrients (e.g. Fe) to the Southern Ocean promoting higher algal productivity and atmospheric CO2 drawdown. Most of the variation in total grounded ice volume in Antarctic glacial-cycle ice sheet models is due to expansion and contraction of grounding lines across continental shelves, mostly in the West Antarctic Ice Sheet (WAIS) Ross and Weddell sectors, and is equivalent to B15–20 m of sea level. The East Antarctic Ice Sheet interior responds in the opposite sense, contracting slightly at Northern Hemispheric glacial maxima due to lower model snowfall rates. In general millennial-scale cycles occur in high-resolution Antarctic ice cores, particularly in methane and temperature records. However, they appear lower amplitude and the timing may differ from their Northern Hemisphere counterparts (Dansgaard–Oeschger or D-O events). Although uncertainties between the age of the gas (methane) and the ice enclosing it in ice cores makes evaluation of the precise timing of inter-hemispheric, millennial-scale cycles difﬁcult, some D-O cycles appear to be synchronous. The pattern of millennial-scale variability superimposed on G-I climate cycles

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in Antarctic ice cores occurs both during and prior to the last 100 kyr cycle. Millennial-scale Antarctic warm periods (A1–A7) are well expressed in SST, sortable silt and d18O records from the southern mid-latitude ocean. One of these events, the Antarctic Cold Reversal, at B14.2–12.4 ka, is associated with an abrupt cooling of B21C in the Southern Ocean, an expansion of ice shelves and sea ice, modest intensiﬁcation of winds and intensiﬁcation in deep abyssal inﬂow along eastern New Zealand through the Paciﬁc gateway.

11.1. Introduction Two major scientiﬁc breakthroughs occurred in the late twentieth century that revolutionised our understanding of Earth’s Quaternary climate system and provided new insight into its pattern of behaviour at both orbital and suborbital time scales. (1) In the mid-1970s, the ﬁrst oxygen isotope records, developed from deepocean sedimentary archives (e.g. Shackleton and Opdyke, 1976), demonstrated that numerous glaciations of the Northern Hemisphere (NH) continents, and associated ﬂuctuations in global sea level, had occurred in response to variations in Earth’s orbital geometry (Hays et al., 1976). Canonical Milankovitch theory (Milankovitch, 1941) had predicted that long-term variations in NH summer insolation inﬂuenced by precession (20 kyr cycle) and obliquity (40 kyr cycle) controlled the variability of the NH ice sheets. Subsequent studies indicated that these orbital inﬂuences produced globally synchronous responses in: (i) atmospheric circulation (e.g. Petit et al., 1999), (ii) ocean circulation (e.g. Rahmstorf, 2002), (iii) terrestrial climate (e.g. Leroy and Dupont, 1994; Ding et al., 2002), (iv) sea level (e.g. Chappell and Shackleton, 1986; Naish et al., 1998) and (iv) Antarctic ice volume (e.g. Denton et al., 1986). (2) In the late 1980s and early 1990s, the ﬁrst deep ice core records from Greenland identiﬁed a pervasive B1,500 yr long climate cycle manifested as a series of warm phases termed Dansgaard–Oeschger (D-O) interstadials that punctuated the otherwise cold conditions of the glacials (Dansgaard et al., 1993; Hammer et al., 1997). These D-O climate oscillations were also identiﬁed in ocean temperature and ice-rafted debris records from NH high-latitude marine sediment cores (e.g. Bond et al., 1997), and indicated a major reorganisation of the NH climate involving switching of the oceanatmospheric system between two principal modes, warm and cold (e.g. Broecker et al., 1990), over decades to just a few years (Severinghaus et al., 1998).

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Millennial-scale climate variability has been associated with far-reaching, but not necessarily globally synchronous inﬂuences. These include: (i) abrupt changes in meridional oceanic overturning (e.g. Broecker, 2000; Van Kreveld et al., 2000; Weaver et al, 2003), ﬂuctuating NH ice volumes (Alley & MacAyeal, 1994) associated with metre-scale changes in sea level (Chappell et al., 1996), atmospheric circulation as recorded by Antarctic ice cores (Blunier et al., 1998; EPICA Community Members, 2006) and the response of Southern Hemisphere (SH) glaciers (e.g. Denton and Hendy, 1994). This burgeoning Quaternary climate dataset, developed largely from NH records, highlighted the dearth of information from Antarctica and the Southern Ocean. A number of important hypotheses and questions began to form concerning the role of Antarctic Ice Sheets and sea ice on ocean circulation and sea level in the Late Pliocene–Pleistocene bipolar glacial world. These included: 1. What is the fundamental response of the Antarctic Ice Sheets to orbital variations? 2. What is the inﬂuence of local insolation forcing, and what is the precise phase relationship between NH and SH climatic processes? 3. Did the dramatic millennial-scale, D-O climate cycles of the NH manifest themselves in the SH? 4. If 3 were true, what was the amplitude of the response, and was the SH in-phase or out-of-phase with the D-O cycles? The short duration of these cycles and the inherent uncertainties with age models of the climate proxy records (marine and ice core) being compared is problematic. 5. To what extent did Antarctic ice volume variability at both orbital and suborbital scale modulate or even drive global climate? This question is especially relevant in terms of its inﬂuence on seasonal sea-ice distribution, the production of deep water and thermohaline circulation. 6. What has been the contribution of Antarctic ice to global ice volume and sea-level ﬂuctuations through the Quaternary? 7. How have the Antarctic Ice Sheets contributed, and evolved, through major Late Neogene climate thresholds such as: (i) the global cooling B3.0– 2.5 myr associated with the initiation of NH glaciation, and (ii) the onset of the 100 ka climate cycle during the Mid-Pleistocene Transition (MPT). Owing to Antarctica’s remoteness, its extensive sea-ice apron and its extended ice sheet cover for the last 34 million years, such seemingly fundamental questions are only now being addressed, through the recovery of well-dated, deep marine sedimentary and ice core archives that span many glacial–interglacial (G-I) cycles. These records largely stem from three major

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international drilling initiatives: (1) the European Programme for Ice Coring in Antarctica (EPICA) and its predecessor project at Vostok station, (2) the Integrated Ocean Drilling Programme (IODP) and its predecessor Ocean Drilling Programme (ODP), and (3) continental margin drilling programmes from ﬂoating ice platforms such as the Dry Valleys Drilling Project (DVDP), Cenozoic Investigations in Ross Sea programme (CIROS), the Cape Roberts Project and the ANDRILL Program. In this chapter we summarise the present state of knowledge of Late Pliocene–Pleistocene Antarctic-Southern Ocean climate behaviour using data and interpretations from Antarctic ice cores (e.g. Vostok and EPICA), Southern Ocean sediment cores (ODP–IODP) and Antarctic continental margin geological drill cores (CIROS-2, Cape Roberts Project, ANDRILL). We also review the advances provided from numerical ice sheet and global climate models, with special emphasis on the potential of the new datasets to be integrated with a new generation of numerical models.

11.2. Glacial Variability from the Continental Margin Geological Record Continuous sedimentary records of Late Pliocene–Pleistocene (last B2.5 myr) glacial variability are generally lacking on the Antarctic continental shelf (deﬁned as the glacially, over-deepened sea ﬂoor above the 1,000 m isobath, Fig. 11.1A). This is because: (1) ice grounding during Quaternary ice sheet expansions onto the shelf has produced signiﬁcant hiatuses, and (2) the ‘riser-less’ rotary coring systems employed by the ODP, and its predecessors, usually result in poor core recovery (o50%) of unconsolidated coarse-grained glacimarine sediments. While, the ﬁnergrained pelagic and hemipelagic sedimentary records recovered from deepwater sediment drifts on the continental rise have yielded more continuous records, these do not provide direct physical evidence of past glacial ﬂuctuations. The history of Antarctic margin Cenozoic geological drilling is summarised by Barrett (this volume, Chapter 3). Here we review in more detail the Late Cenozoic intervals of those records. Since the mid-1970s, ship-based continental margin drilling of Late Pliocene–Pleistocene sediments has been focused in three main regions (Fig. 11.1): (1) Antarctic Peninsula (ODP Leg 178; Barker et al., 1999) and Weddell Sea (ODP Leg 113; Kennett and Barker, 1990); (2) Prydz Bay (ODP Leg 119; Barron et al., 1989; ODP Leg 188; O’Brien et al., 2001); (3) Ross Sea (DSDP Leg 28; Hayes and Frakes et al., 1975; Hays et al., 1976). Average core recovery on the continental shelf has ranged from 8%

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(Leg 178 sites 1097, 1100, 1102, 1103) to 30% (Leg 119, sites 739–743) (see Fig. 11.1 and this volume, Chapter 3). The continental margin off Wilkes Land is scheduled for drilling by the IODP in January 2009 (IODP expedition 318, Escutia et al., 2006). In parallel with ocean drilling activities, a different drilling approach based on a closed-circulation, continuous wireline coring system used by the minerals industry, was being applied from land and sea-ice platforms in the McMurdo Sound region. The ﬁrst glacimarine Pleistocene records were recovered by the DVDP (McGinnis, 1981) from the deglaciated mouth of the Taylor Valley with signiﬁcantly better core recovery (B90%) than ocean drilling. The ﬁrst attempts at drilling from sea ice were made in the late 1970s with limited success and core recoveries of 50–60% (DVDP-15; Barrett and Treves, 1981; MSSTS-1; Barrett, 1986). During the 1980s and 1990s progressively more sophisticated sea-riser technology was employed by the CIROS (Barrett, 1989) and the Cape Roberts Project (CRP; CRST, 1999) that enabled the ﬁrst deep Cenozoic records to be recovered (95–98%). Nevertheless, Pliocene–Pleistocene strata remained poorly sampled due to glacial erosion and non-deposition at these sites. A region that had experienced high rates of Pliocene–Pleistocene basin subsidence and sedimentation was identiﬁed beneath the McMurdo Ice Shelf in the centre of the Victoria Land Basin (Naish et al., 2006), and was drilled by the ANDRILL Programme’s McMurdo Ice Shelf (MIS) Project in the austral summer of 2006–2007 (Naish et al., 2007). The 1,285 m long AND-1B drill core recovered by the MIS Project provides the most complete (98% core recovery) and continuous sedimentary record of Pliocene–Pleistocene climate and glacial history from the Antarctic continental margin to date (Naish et al., 2008) and is reviewed in more detail below.

11.2.1. Weddell Sea Pliocene–Pleistocene sediments were recovered in the Weddell Sea during ODP Leg 113 from three distinctly different deep-water environments Figure 11.1: (A) Location of ice cores from the Antarctic interior and ocean drill cores (DSDP, IODP, ODP) from the continental margin and rise discussed in text. (B) Location of sea-ice and ice-shelf based geological drill core records from the Victoria Land Basin (VLB) in McMurdo Sound. (C) Geological cross section (A-A’, Fig. 11.1B) across the VLB shows the stratigraphic relationships and age of strata recovered by drill cores in Southern McMurdo Sound.

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(Fig. 11.1): hemipelagic and biogenic sedimentation on the South Orkney microcontinent inﬂuenced by the WAIS (Site 697), a turbiditic sequence in the deep Weddell Basin inﬂuenced by both WAIS and the East Antarctic Ice Sheet (EAIS) (Site 694) and hemipelagic and terrigenous sedimentation on the east Antarctic continental margin (site 693). From these sites, Kennett and Barker (1990) were the ﬁrst to identify a cooling step during the Late Pliocene that they associated with development of the present cold polar thermal regime of the Antarctic Ice Sheets. This is marked between 3 and 2.5 Ma, by a regional reduction in sedimentation rates and diatom abundances, a deterioration in microfossil preservation and a dominance of the sea-ice diatom Eucampia antarctica. At this time cooling led to expansion of WAIS and EAIS to the edge of the continental shelf coincident with the development of the NH continental ice sheets. Leg 113 cores also showed that Quaternary expansions of the Antarctic Ice Sheet displayed pronounced orbital-scale, G-I sedimentary cyclicity characterised by relatively low siliceous biogenic components and higher average particle size during glacial deposition, alternating with larger silt-sized grains and higher diatom abundances during interglacials. Thus, glacial periods were characterised by increased sea-ice cover, lower productivity and stronger bottom water ﬂows (Pudsey et al., 1988; Pudesy, 1990). The converse was true for interglacial periods. With the exception of the G-I cycles in the upper 20 m of Site 697 core, which were correlated with Marine Isotope Stages 10-1, the chronostratigraphy did not allow accurate dating of the Early and Middle Pleistocene, G-I cycles. An enigmatic, 2.5 m thick interval of foraminiferal ooze, bearing coccolith assemblages occurred in a number of Leg 113 cores, and implied an interval of warming and depression of the lysocline. While the age of this calcareous unit remains poorly constrained in the Weddell Sea region, it may correlate with a similar unit reported from Prydz Bay and Ross Sea sediment cores and dated at B1 Ma (see below). 11.2.2. Antarctic Peninsula Poor core recovery from shelf sites during Leg 178 drilling on the Paciﬁc margin of the Antarctic Peninsula resulted in virtually no sedimentary record (Barker et al., 1999) of the progradational wedges that are inferred to have built the continental shelf during Quaternary oscillations of the WAIS (Larter and Barker, 1989; Larter and Cunningham, 1993). Notwithstanding this, the characteristics of the Pleistocene diamicts and glacimarine sediments that were recovered from the shelf sites were consistent with sedimentation

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on the margin since the Last Glacial Maximum (LGM). This observation led Barker et al. (1999) to propose that the topset deposits known as Antarctic Margin seismic sequence unit S1 were deposited under the inﬂuence of an expanded and periodically grounded WAIS. Rebesco et al. (2006) correlated a grid of seismic reﬂection data collected on the Paciﬁc shelf margin after Leg 178 into the more complete and betterdated, continental rise Site 1101. They showed a transition across the seismic sequence unit S2-S1 boundary marked by a change from deep erosion and debris ﬂows to the progradation of relatively stable wedges on the continental slope occurred B3.0 Ma. The timing is coincident with a signiﬁcant reduction in sedimentation rate over the distal margin area consistent with observations on the other side of the Antarctic Peninsula in the Weddell Sea (Kennett and Barker, 1990). Both studies argued that the change in depositional style corresponds to major ice sheet expansion onto the continental shelf associated with reduced sediment supply and meltwater volume as the Antarctic Ice Sheets attained their present cold polar state. Drilling of continental rise sediment drifts at sites 1095 and 1096 produced a more continuous record of Pleistocene G-I variability expressed as lithologic alternations between interglacial biosiliceous silts and glacial nonfossiliferous, terrigenous laminated clays and silts. Intriguingly, time-series analysis of the properties of G-I cycles at sites 1095 and 1096 shows no periodicity in the orbital frequency bands. However, iceberg-rafted debris (IRD) mass accumulation rates in the Pliocene–Pleistocene drift record at Site 1101 show strong covariance with the B40 kyr, obliquity frequency between 3 and 1 Ma, after which 100 kyr frequencies dominate (Cowan, 2002). Two large spikes in IRD abundance at 2.8 and 0.88 Ma have been interpreted as corresponding to Late Pliocene cooling and ice sheet expansion onto the shelf, and a second phase of cooling and ice expansion across the MPT (see below). The occurrence of orbitally inﬂuenced ice rafting in records around the Antarctic Peninsula from B2.8 Ma has led a number of workers to suggest a NH eustatic inﬂuence on the Quaternary stability of the Antarctic Ice Sheet margin (e.g. Barker et al., 1999; Cowan, 2002; Hillenbrand and Ehrmann, 2005). 11.2.3. Prydz Bay Drilling was undertaken at ﬁve sites (739–743, Fig. 11.1) during ODP Leg 119 on a transect across the continental shelf of Prydz Bay to elucidate the long-term glacial history and evolution of the EAIS (Barron et al., 1989). Although the late Pliocene to Pleistocene interval is limited by poor core

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recovery and inadequate age control, a cyclic record of tills and ice proximal, glacimarine sediments implies a succession of advances and retreats by the EAIS across the region (Hambrey et al., 1991). During Leg 188, two sites sampled Late Pliocene–Pleistocene strata (O’Brien et al., 2001). Site 1167 was the ﬁrst hole to be drilled in an Antarctic margin trough mouth fan, and the core provides an expanded view of glacier ﬂuctuations in the Prydz Bay region during the last 2 Ma. The sedimentary record is dominated by several metres thick diamictons deposited as debrites separated by thin mud units (O’Brien et al., 2007). The debrites represent slumping and remobilisation of glacial outwash deposits and tills when the EAIS grounding line is near the shelf edge. Intervening interglacials are represented by glacimarine muds deposited by the retreating ice sheet. While the diamicton-mud cycles maybe of orbital duration, hiatuses within the section and insufﬁcient age control preclude the duration of cycles from being established. O’Brien et al. (2007) also pointed out that the trend to expanded ice in the Prydz Bay region during the Late Pliocene maybe associated with the movement of interior precipitation to the coast, either from changes in storm tracks or the increase in ice around the EAIS margin. Moreover, they argue that this process could have caused the change to more laterally conﬁned ice streams since the early Pliocene. A signiﬁcant change in sedimentation rate and sediment composition suggests a colder and less erosive ice sheet developed following the MPT at about 0.78 Ma. Three of the sedimentary cycles above the Bruhnes– Matuyama (B-M) magnetic reversal boundary (0.78 Ma) correspond to three cycles in a d18O record from Neogloboquadrina pachyderma at Site 1167 that are tentatively correlated with Marine Isotope Stages 21-16 and indicate orbital control on ice sheet ﬂuctuations (Thiessen et al., 2003). Site 1165 sampled sediment drift deposits on the continental rise and comprises a strongly cyclic Pleistocene interval in its upper 50 m in particle size, grey-scale and clay mineral records (O’Brien et al., 2007). A recent revision of the Pleistocene bio- and magnetostratigraphy of the sites show an incursion of calcareous nannofossils in the Site 1165 core at around 1 Ma (Villa et al., 2007), coincident with a decrease in d18O values in records from both 1167 and 1165 (Thiessen et al., 2003). Late Pliocene–Early Pleistocene glacimarine deposits are described on land from the uplifted sequences of the Bardin Bluffs Formation (uppermost formation of the Pagodroma Group) in the Prince Charles Mountains (e.g. Whitehead and McKelvey, 2001). While the formations within this group are geographically separated they have been correlated into a composite stratigraphic section on the basis of diatom biostratigraphy (e.g. McKelvey et al., 2001; Whitehead et al., 2004). Bardin Bluffs Formation

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sediments mark a dramatic change from older warmer intervals of diatombearing glacimarine sediments to coarse-grained, ice-proximal diamicts overlying a striated erosion surface above the Amery Group. 11.2.4. Ross Sea The pioneering coring expedition of DSDP Leg 28 in the Ross Sea drilled four sites on the continental shelf: Sites 270–272 in the south eastern quadrant of the Ross continental shelf and Site 273 in the west-central portion of the shelf (Hayes and Frakes et al., 1975) (Fig. 11.1). While core recovery was low (B30%), a composite Pliocene–Pleistocene record of W200 m was sampled at sites 270–272 comprising terrigenous silts and biosiliceous-bearing muds and diatomaceous ooze with ice-rafted debris throughout. A cyclostratigraphic interpretation was not reported for these sites, but a major glacial unconformity of inferred Late Pliocene age identiﬁed in Ross Sea seismic data, and correlated into Site 270, is interpreted as widespread westward expansion of the WAIS ice streams into the Ross Sea (Bart, 2004). Rebesco et al. (2006) have correlated this event with evidence elsewhere for continentwide development of a marinebased ice margin on the continental shelf, associated with development of a thicker, cold ice sheet similar to the present-day thermal regime. Immediately prior to the coring of Leg 28 the Dry Valley Drilling Project (DVDP) recovered the ﬁrst Pliocene–Pleistocene glacimarine sediment cores from the continent (McGinnis, 1981). A B200 m thick composite section based on DVDP Sites 10 and 11 in the Lower Taylor Valley, comprised a succession of ice-proximal Late Pliocene glacimarine diamictons, sands and gravels passing up-section into ﬂuvio-delatic and beach sands of Late Pleistocene age (McKelvey, 1981; Powell, 1981). Lateral facies relationships and correlations between the two sites indicate that the down valley Site 10 was most proximal to the grounding line. This relationship together with the occurrence of McMurdo volcanic clasts in diamictons implied an eastern ice source during glacial advances associated with expansions of the WAIS (Porter and Beget, 1981). McMurdo volcanic-bearing diamictons occur above an unconformity at 203 m depth and 146 m depth in DVDP 11 and 10, respectively, corresponding to the proximity of the Gauss–Gilbert polarity transition (2.6 Ma) (Elston and Bressler, 1981). Thus, although some uncertainties surround the chronology of these cores remain, expansion of WAIS-sourced ice into the Dry Valleys is consistent with timing of Late Pliocene ice expansion elsewhere. Barrett and Hambrey (1992) report a Late Pliocene–Pleistocene succession of McMurdo volcanic-bearing tills in

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the upper 100 m of the CIROS-2 core from Ferrar Fiord. They interpret the succession to represent the onset of ice advance from the Ross/WAIS Ice Sheet at around 3–2.5 Ma. 11.2.5. The ANDRILL Programme: McMurdo Ice Shelf Record of Late Pliocene to Pleistocene Climate Variability During the austral summer of 2006–2007, a new Antarctic geological drilling programme, ANDRILL, successfully recovered a 1,285 m long record of strata (AND-1B) from beneath the McMurdo Ice Shelf sector of the Ross Ice Shelf. This record spans the last 14 myr (Fig. 11.2; Naish et al., 2007).

11.2.5.1. Orbital-scale, glacial–interglacial sedimentary cycles Forty unconformity-bound glacimarine cycles in the upper 600 m of the core record orbitally inﬂuenced ﬂuctuations in the areal extent of the WAIS margin in the Ross Embayment, and the evolution of the whole Antarctic Ice Sheet from a period of relative global warmth in the Early Pliocene through a profound cooling step in deep-sea oxygen isotope records B3–2.5 Ma, to the development of the present cold polar Antarctic thermal regime during the last B1 million years (Naish et al., 2007, 2008). The rocks were interpreted in terms of lithofacies associations – sediments representing speciﬁc environments of deposition. These ranged from open marine diatomites, mudstones and turbidites deposited during interglacials to ice-proximal massive and stratiﬁed diamictites, conglomerates and sandstones representing glacial periods. During glacial periods the ice sheet had a laterally extensive marine terminus extending well out into the Ross Sea beyond the drill site. In interglacials the drill site was covered by either an ice shelf (similar to present day), or lay in open water when the ice sheet margin had retreated onto the continent, with local deposition of marine diatoms, terrigenous mud and occasional debris from ﬂoating ice.

11.2.5.2. Chronology A preliminary age model for the upper 700 m of drill core constructed from diatom biostratigraphy and radiometric ages on volcanic material allows a unique correlation between B70% of the magnetic polarity stratigraphy and the Geomagnetic Polarity Time Scale (GPTS) (Wilson et al., 2007).

Figure 11.2: Lithological column of the upper 600 m of the AND-1B drill core recovered by the ANDRILL MIS Project. Correlation of biostratigraphically constrained palaeomagnetic reversals (Wilson et al., 2007) with the oxygen-isotope time scale (Lisiecki and Raymo, 2005) are shown. Glacial surfaces of erosion mark boundaries of orbital-scale, glacimarine sedimentary cycles (Naish et al., 2008). Density, clast abundance and glacial proximity curves highlight the glacial and sedimentary cyclicity of G-I cycles. I, Ice Contact; P, Proximal Glacimarine; D, Distal Glacimarine; M, Open Marine.

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The age model provides several well-constrained intervals displaying relatively rapid (o1 m/1,000 yr) and continuous accumulation of sediment punctuated by several 0.5–1 myr stratal hiatuses representing more than half of the last 7 myr. Thus, the AND-1B record provides several highly resolved ‘windows’ into the Late Cenozoic development of the Antarctic Ice Sheets.

11.2.5.3. Implications for late Cenozoic ice sheet history The Pliocene period (5–2 Ma) is characterised by a dynamic ice margin with interglacials displaying pelagic diatomite, implying high phytoplankton productivity in locally open water. A W80 m thick interval of diatomite between 370 and 460 m below sea ﬂoor (mbsf) represents an extended period of open water in the Ross Embayment and high phytoplankton productivity. Two erosional unconformities of greater than 100 kyrs dated at B3.3 and 2.7 Ma enclosed by diamictites occur at B285 and 250 mbsf, and are interpreted to represent stepwise cooling, expansion of the WAIS and glacial erosion in Western Ross Embayment. The timing leads and is coeval with the development of the NH ice sheets and evidence for ice expansion on other regions of the Antarctic continental shelf already discussed. (e.g. Rebesco et al., 2006). Stratigraphic analysis and a new high-resolution integrated chronostratigraphy (Wilson et al., 2007) for the overlying Late Pliocene AND-1B cycles, particularly the abrupt transitions from diamictites to biosiliceous deposits, implies rapid oscillations between glacial and marine environments with signiﬁcant volume changes in the WAIS at a frequency of 41 kyr. Intriguingly, the facies also imply a cooler style of ice sheet compared with the Early Pliocene cycles, indicating reduced amounts of subglacial meltwater and local terrigenous sediment input from the McMurdo Sound area (McKay et al., accepted). Naish et al (submitted) propose that between 3.5 and 2.5 Ma, high-latitude climate cooling drove both the WAIS and EAIS towards their present expanded cold polar state. Relatively warm, often terrestrially based ice margins were replaced by more permanent marine termini and the development of ice shelves. Mass balance changes affecting these colder ice sheets may have occurred primarily through calving processes at the marine margins, rather than melting. From B2.6 Ma polar Antarctic ice volume changes may have became signiﬁcantly inﬂuenced by NH, glacio-eustasy and this mechanism may account for much of the orbital variability of the WAIS since B2.6 Ma.

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A dramatic change in sedimentary cycle architecture occurs above an erosion surface at 82.74 mbsf in the AND-1B core, that is more or less coincident with the B-M polarity transition, 0.78 ka. Eight ﬂuctuations in the proximity of WAIS margin, expressed as, polar-style sedimentary cycles, occur within the interval above the B-M boundary and correspond to an interval of eight, 100 kyr duration d18O oscillations between Marine Isotope Stages 20 and 2. These diamictite-dominated cycles indicate a predominance of a cold polar ice sheet over the drill site for most of the Late Pleistocene, which retreated to form an ice shelf during interglacial minima. The erosional surface beneath these cycles removes about B100 kyr of the upper part of the Matuyama Chron above a normal polarity interval of biosiliceous-bearing mudstone assigned to the Jaramillo Subchron (Wilson et al., 2007). An anorthoclase-bearing tuff within the mudstone has yielded a 40Ar/39Ar age of 1.01470.004 Ma (Wilson et al., 2007) and indicates that this period of open marine pelagic sedimentation in western Ross Sea occurred during the warm interglacial of Marine Isotope Stage 31 providing the last evidence for an open Ross Embayment in the AND-1B record. The calving line of the Ross Ice Shelf appears not to have retreated south of its present interglacial position during subsequent ‘warmer-than-Holocene super-interglacials’ (e.g. Marine Isotope Stages 11, 9 and 5; Jouzel et al., 2007). These observations of Late Pleistocene relative stability of the Ross Ice Shelf seem incompatible with the geological evidence for Late Pleistocene collapse of the WAIS (e.g. Scherer et al., 1998), and remnant shorelines 20 m above present sea level during these interglacials (Hearty et al., 1999).

11.3. Atmospheric Variability from Ice Cores So far ice cores reach only through the last third (800 ka) of the Pleistocene epoch (Jouzel et al., 2007). However, for that period they provide iconic datasets that can help to unlock some aspects of Antarctic climate evolution. This is because they supply a record of Antarctic climate (temperature and precipitation) that drives the evolution of the ice sheet, and they also supply the records of the most important forcings (especially greenhouse gases, but less directly other forcings); these have in turn presumably controlled the climate. One thing ice cores cannot constrain so well is the response of the ice sheets to climate change, and that is the value of the continental shelf and deep marine geological records, and the integration of both.

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11.3.1. The Nature of the Ice Core Record Ice cores are a sedimentary record like many others, but comprise snow and ice rather than terrigenous particles, with layers deposited sequentially, although unlike other records the layers thin with depth due to ice ﬂow. They have the beneﬁt that they record the atmosphere (including sampling precipitation, aerosol and trace gases) more directly than any other archive, without biological mediation. Highly resolved records, potentially datable by counting annual layers, are recovered at sites with high snow accumulation rates. The longest records extend back hundreds of thousands of years. However, at these sites annual layers cannot be resolved and far more difﬁcult exercises are needed in order to derive a reliable age scale (e.g. Parrenin et al., 2007a). A particular strength of ice cores is that so many aspects of climate forcing and response are recorded in one core, while a weakness (less obvious for this book that concentrates on Antarctic climate) is that most ice core records come from the polar regions (particularly Greenland and Antarctica). The ice core record is held in three distinct forms. Firstly, the water molecules themselves, through their isotopic content, contain a record of the 16 temperature at the time the snow fell. The ratio of H18 2 O/H2 O, or of HDO/ 2 H2O (where D is deuterium, H) is generally considered to be a good proxy for site temperature, and the primary record of most ice core studies. This arises from the fact that, after water vapour has been evaporated from a warm ocean, it steadily loses water as it cools on the way to the poles, with the heavier isotopes having a lower vapour pressure and, therefore, condensing more readily, which leaves an airmass depleted in heavy isotopes to continue towards the poles. Although the real relationship is of course more complicated (Jouzel et al., 2003), in essence this leads to more negative isotopic values (expressed as per mil deviations from the composition of ocean water) in colder climates. As an aside, it is worth noting that it is the buildup of ice with negative isotope values in the ice sheets that leads to positive isotopic excursions in the oceans during colder (more glaciated times). When snow falls, it takes with it aerosol particles and some gases that have an afﬁnity for water. More are added just after deposition. It is in this part of the record, of material trapped in surface snow, that we ﬁnd for example the spikes of sulphuric acid that are found in snow for a few years after major volcanic eruptions, as well as components of sea salt and terrestrial dust. Finally, snow turns to solid ice in the cold polar regions only under the pressure of overlying ice that sinters the snow grains together. This is the process that seals samples of atmospheric air in air bubbles, typically at

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60–100 m depth, and it is in these bubbles that all the stable trace gases, including carbon dioxide and methane, are found. As already discussed, most ice cores come from the polar regions. In Greenland, the oldest continuous ice core so far is from North GRIP (North Greenland Ice Core Project Members, 2004), at 123 kyr, reaching just into the last interglacial. Several other cores, spanning the latitudes of Greenland, cover a slightly shorter period continuously. In Antarctica (Fig. 11.1), the oldest ice cores, spanning several G-I cycles, come from the East Antarctic plateau sites with very low snow accumulation rates. For many years, Vostok, spanning four glacial cycles and 420 kyr (Petit et al., 1999), was the oldest ice core, and it remains the longest in terms of depth. A slightly shorter period has so far been published from Dome Fuji (Watanabe et al., 2003), although a new core has recently extended the age span towards that of Dome C. The European EPICA ice core at Dome C now contains the longest (in age) continuous sequence, at just over 800 kyr (EPICA Community Members, 2004; Jouzel et al., 2007). A number of other cores have been drilled in the coastal regions of Antarctica that reach beyond the LGM (e.g. Law Dome, Siple Dome, Berkner Island); these are important because they hold the potential to derive records of climate and ice sheet size in the outlet regions of the continent. Finally, central West Antarctica is so far represented only by the Byrd core, but a new effort from the US Antarctic Programme (WAIS Divide Project) promises a new high-quality core in the next few years.

11.3.2. Antarctic Climate and Forcing at Glacial–Interglacial Time scales Rather than give a historical account of increasing knowledge in this area (which would focus strongly on Vostok), we will here use the longest (in age) record from Dome C to describe what is known, and then discuss what the similarities and differences of other records indicate. Figure 11.3 shows the record (using water isotopes) of Antarctic temperature at Dome C over the last 800 ka (Jouzel et al., 2007). The age scale has been derived using an ice ﬂow and snow accumulation model constrained, and in some parts tuned, to independent age markers of known age (Parrenin et al., 2007a). The maximum uncertainty of the absolute ages is estimated to be better than 6 kyr throughout the core. Through most of its span it shows good coherence with signals recorded in the benthic marine oxygen isotope stack (Fig. 11.3, representing some combination of global ice

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Figure 11.3: Top panel: marine benthic oxygen isotope stack on LR04 age scale (Lisiecki and Raymo, 2005). Lower panels: 800 kyr record, on EDC3 age scale of Deuterium, and temperature (difference from the last millennium, derived using a correction for seawater isotopic content and modelled ice-sheet altitude) (Jouzel et al., 2007). Numbers above the marine curve represent selected Marine Isotope Stages. volume and deep-water temperature) whose time scale was derived in a completely different way (Lisiecki and Raymo, 2005). The ﬁrst obvious feature of the record is that a generally cold climate (up to 91C colder than the late Holocene in this part of Antarctica) was interrupted approximately every 100 kyr by warmings (such as the Holocene interglacial) lasting approximately 10–30 kyr. The last four interglacials (back to Marine Isotope Stage 11) had periods that were 2–4.51C warmer than the last millennium. Before Marine Isotope Stage 11, a different pattern is seen: although 100 kyr still emerges as the dominant period of the record (Jouzel et al., 2007), the earlier interglacials are signiﬁcantly cooler than the later ones, and the system seems to spend a higher proportion of time in the warmer phase. There is not yet any convincing explanation for this change in amplitude of the signal (discussed below). The next question that arises is whether the Dome C record can be considered representative for a larger part of the Antarctic. Comparison with

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Vostok (around 500 km away) shows excellent agreement, but more importantly a very similar signal (albeit with small differences in the age scale) is seen at Dome Fuji, on the opposite side of the East Antarctic plateau (Watanabe et al., 2003; Kawamura et al., 2007). This similarity suggests that the Dome C record can stand as a signal of climate across much of the polar plateau. The 100 kyr periodicity seen in the ice core record is of course already well-known from other climate records, and assumed to derive in some way from orbital forcing ampliﬁed by internal feedbacks. One of the most important ampliﬁers is the greenhouse gas concentration of the atmosphere. The records of greenhouse gases from Dome C (completed by those from Vostok in as yet unmeasured sections of Dome C) extend so far to 650 ka (and are currently being continued to 800 ka). They show very strong congruence with many features of the temperature record, and are consistent with CO2 in particular playing a signiﬁcant role in such ampliﬁcation (Figs. 11.3 and 11.4). Concentrations of CO2 are typically around 180–200 ppmv in the coldest parts of each cycle, and reached 280–300 ppmv during the warm periods back to Marine Isotope Stage 11 (Siegenthaler et al., 2005). They were however considerably lower (B250 ppmv) during the warm periods before Marine Isotope Stage 11, scaling with Antarctic temperature (Fig. 11.4). The very strong similarity of CO2 concentrations and Antarctic temperature strongly suggests that the Southern Ocean plays a leading role in controlling the atmospheric concentration of CO2 on these time scales, as already implied by more detailed records of the last deglaciation (Monnin et al., 2001). Concentration of CO2 in the atmosphere today (ca. 380 ppmv) is nearly 30% higher than the highest value seen in the previous 650 kyr. Methane (CH4) concentrations (Spahni et al., 2005) show strong similarities to the Antarctic temperature record (Fig. 11.4), with high concentrations in warm periods, low concentrations in glacial maxima and lower concentrations in ‘weaker’ interglacials compared to ‘stronger’ ones. However in this case, even at the low resolution shown, it is apparent that there are higher-frequency millennial-scale variations superimposed on the orbital-scale trends. In the most recent climatic cycle (Blunier and Brook, 2001) these correspond to the D-O events that will be discussed later. Methane concentrations are most likely mainly controlled by changing wetland extents in the northern high latitudes and tropics, along with changes in sinks (Valdes et al., 2005), and thus are relevant to Antarctic climate only as a relatively minor ampliﬁer. Ice cores have provided other data that are important for understanding Antarctic climate and ice sheets. The snow accumulation rate is clearly

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Figure 11.4: Records of Deuterium, non-sea-salt Calcium ﬂux (Wolff et al., 2006), CO2 (Siegenthaler et al., 2005) and CH4 (Spahni et al., 2005) measured at Dome C. Note that there maybe small timing mismatches because the gases are shown on the EDC2 age scale, while the other records are on EDC3. important for the mass balance of the ice sheet. In central Antarctica, it has generally been treated by a modelling approach in which the accumulation rate is thermodynamically related to the local temperature. Snow accumulation scales non-linearly with temperature, and was around half its present value in central East Antarctica during the coldest periods of the last 800 kyr than it is at present. Ice sheet elevation in central Antarctica is assumed to have varied little over the last 800 kyr: simple modelling exercises (Parrenin et al., 2007b) used as input to age models suggest variations of less than 200 m at both Dome C and Dome Fuji. Measurements of air volume (which has atmospheric pressure as one among its controlling variables) provide only a limited constraint, conﬁrming that altitude changes of more than a few

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hundred metres have not occurred. Of course, it is possible for the areal extent of an ice sheet to change considerably without any great change in the altitude of its central parts. The dust ﬂux to Antarctica (represented in Fig. 11.4 by non-sea-salt calcium) was much higher during glacial maxima than during warm periods. This is believed to be a result of changes in climate in southern South America (Wolff et al., 2006), although the exact nature of the relevant changes is not yet clear. In any case the result is a change in aerosol content of the atmosphere over Antarctica (with possible radiative impacts) and a change in the amount of dust deposited in the Southern Ocean which may also play a role in stimulating phytoplankton productivity, and thus carbon dioxide drawdown from the atmosphere. Finally in this section, it has been suggested that the sea salt content of Antarctic ice cores might yield a record of past sea ice extent (Wolff et al., 2006). Sea salt, as expected under this idea, is at higher concentrations during cold periods, consistent with the ﬁndings of marine sediments (Gersonde et al., 2005) that sea ice was signiﬁcantly extended at such times. 11.3.3. Millennial-Scale Changes in Antarctic Climate The climate of the last glacial period is strongly dominated by the occurrence of rapid millennial-scale climate changes known as D-O events (e.g. North Greenland Ice Core Project Members, 2004). These occur in Greenland as rapid (occurring in o40 years) warmings by more than 101C (Huber et al., 2006). Counterparts to them are clearly seen in numerous other NH records. What concerns us here is what was happening in Antarctica in relation to those events, and what it tells us about the causes. The longest D-O events have very clear counterparts, initially labelled as ‘A-events’ in Antarctic climate records. Using the high time-resolution available from the Byrd ice core, Blunier and Brook (2001) showed that while Greenland was cold before a D-O warming, Antarctica experienced a steady warming (by up to B31C). As soon as Greenland temperature jumped, Antarctic temperature started to cool again. It was suggested that this was the pattern expected from simple models involving changes in ocean heat transport (Stocker and Johnsen, 2003), with Antarctica and the Southern Ocean able to accumulate heat when ocean heat transport to the north was weak. Further support for this model has come from the ﬁnding (using the EPICA cores at both Dome C and Dronning Maud Land (Kohnen station)) that there is a subdued southern counterpart to every D-O event (EPICA

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Community Members, 2006). It has not yet been possible to show whether the timing of the weaker, shorter events is consistent with the model. However, central Antarctic climate of the last glacial appears to have been dominated by slow warmings and coolings of between 0.5 and 31C that are most likely associated with changes in meridional ocean heat transport, and thus with rapid warmings in the north. The southern warm events have been named as ‘Antarctic Isotopic Maxima (AIM)’ (Blunier and Brook, 2001), with the A-events representing the largest AIM. Very similar temperature variability was found in the Dome C record for previous glacial periods as for the most recent one, strongly suggesting that the same mode of millennial variability held sway for the last 800 kyr (Jouzel et al., 2007).

11.3.4. Climate of the Antarctic Periphery Only a small number of ice core records from the Antarctic Ice Sheet periphery extend beyond the LGM. The 1,178 m deep ice core at Law Dome appears to reach into the last interglacial period (Morgan et al., 1997), although the whole of the last glacial period is held within the lowest 100 m, and no attempt has been made to provide an age scale for the oldest part. At least during the last deglaciation, the timing of the signal from this core appears at least similar to that from other Antarctic cores. Of most interest for the current paper is the ﬁnding that no isotopic values typical of inland ice are found anywhere in the core. This implies that, although the ice at Law Dome may have thickened during some parts of the glacial by up to 300 m (Delmotte et al., 1999), it was never overridden by inland ice during the last 100 ka. Working clockwise around the continent, Taylor Dome (TD), adjacent to many of the marine core records, has yielded a 554 m deep core to bedrock (Steig et al., 1998). The age scale has been tentatively extended to 300 ka (Grootes et al., 2001); however, Termination II (130 ka) is only 23 m above bedrock, so the record is most useful only for one climatic cycle. Comparison with Vostok during the last climatic cycle shows a very similar pattern of variations, but with a larger amplitude for the equivalents of the A-events (large AIM). The assumptions made to create the time scale for the glacial part of the record preclude any discussion about the phasing of these changes in the TD record relative to those of central Antarctica. There has of course been a suggestion (Steig et al., 1998) that, at least during the last termination, TD may have shown a different pattern of change from other Antarctic sites, more in line with those of Greenland. This assertion has been strongly

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challenged (Mulvaney et al., 2000) because of issues about the TD time scale in this section, and must be considered doubtful. The 1,004 m deep bedrock core at Siple Dome (SD), on the West Antarctic side of the Ross Sea, also covers much of the last glacial period. Careful synchronisation of this core to that of Byrd and to Greenland cores (Brook et al., 2005) shows a similar pattern and timing of millennial-scale change to that of Byrd, with recognisable A-events, and with SD (like Byrd) warming while Greenland remained cold, and apparently starting to cool when Greenland warmed. An abrupt climate change, not observed in other Antarctic cores, is seen at SD at 22 ka (Taylor et al., 2004); there is as yet no clear explanation for this apparently very local climate shift. Two other near-coastal cores have already penetrated beyond the LGM, at Talos Dome and Berkner Island, although publications from them are still awaited. Early indications from Berkner Island (facing the Weddell Sea) suggest that it was also never overridden by inland ice during the last climatic cycle (Mulvaney et al., 2007) as some models suggest. Taking all the records, from central East Antarctica, coastal cores and Byrd in central West Antarctica, it appears likely (despite the questions raised by the TD core) that the continent sees the main millennial- and orbital-scale changes more or less synchronously and with a similar pattern. Further work to separate changes in ice sheet altitude from changes in climate will elucidate the causes of somewhat different amplitudes for millennial-scale change at different sites. Finally in this section, we note the potential for coastal cores such as Law Dome and Berkner Island to provide critical constraints on the extent of the Antarctic Ice Sheet during at least the last glacial cycle. 11.3.5. Future Challenges for Ice Coring The ice core community has prepared a strategy for the next decade or more that includes two tasks of particular importance for this chapter (Brook and Wolff, 2006). Firstly, the community has agreed on the value of further cores around Antarctica (and the Arctic) reaching at least 40 ka back in time, to further understand the spatial and temporal pattern of millennial-scale changes. Secondly, sites with continuous ice sequences reaching well beyond 800 ka should be targeted for drilling. While a major geophysical survey will be required to locate such ice, the goal of reaching back as far as 1.5 Ma, into the period when 40 kyr cyclicity is well-established in climate records, is of great importance. It offers the possibility for example to test ideas about the role of changing CO2 concentrations, and the relative roles of

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northern and southern ice (Raymo et al., 2006), in the transition from the 40–100 kyr world.

11.4. Oceanic Variability from Southern Ocean Sediment Cores As Antarctica changes, so does the ocean. The long-term evolution of the ice sheets and the shorter-term oscillations superimposed upon them, have produced far-reaching effects on the circulation and water masses of both hemispheres. In this section, we focus on the oceans’ responses to G-I cycles over the last million years with emphasis on the last few cycles simply because they are generally the best documented. This younger part of the record is also assessed for the effects of millennial-scale change as driven from Antarctica but also from the NH. Here we review brieﬂy the wealth of palaeoceanographic research and the reader is referred to the reference list as a window into an extensive literature.

11.4.1. Glacial–Interglacial Cycles Past inﬂuences of Antarctica on the Southern Ocean are best gauged from well-dated records of environmental proxies, which are interpreted against a history of Antarctic ice behaviour preserved in ice cores and sediment cores from high-latitude locations. While we focus on Antarctica, it is made with full acknowledgement of NH forcing especially as the dominant driver of eustatic sea-level oscillations and as a co-driver of thermohaline circulation.

11.4.1.1. Thermohaline circulation Essentially, the sinking of upper ocean waters in northern high latitudes forms North Atlantic Deep Water (NADW) that migrates south to mix with southern-sourced waters comprised of (i) dense bottom waters generated around the Antarctic margin and (ii) deep waters recycled in the Indian and Paciﬁc oceans especially within the Antarctic Circumpolar Current (see Carter et al., this volume Chapter 4). The resultant Circumpolar Deep Water is the most voluminous in the Southern Ocean and dominates the deep inﬂows into the Indian and Paciﬁc basins (Fig. 11.5). Yet despite its hybrid character, it retains the signatures of its source waters in particular NADW, which is identiﬁed by the salinity maximum and, from a palaeoceanographic

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Figure 11.5: Generalised outline of the modern Southern Ocean with the westerly wind-driven Antarctic Circumpolar Current deﬁned by its zonal jets or fronts that include the Subantarctic Front (SAF), Polar Front (PF), Southern Front (SF) and Southern Boundary (SB) after Orsi et al. (1995). ODP Sites 1090 and 1123 containing W1 myr records of ocean change are also located.

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perspective, by its elevated d13C content (Charles and Fairbanks, 1992). Changes in this circulation over G-I cycles are subject to a continuing debate, much of which focuses on the input of NADW to the Southern Ocean. Charles and Fairbanks (1992), Rahmstorf (2002), Ninnemann and Charles (2002) amongst others favour a reduction in NADW input to the Southern Ocean in glacial times. Hall et al. (2001) support this decline and suggest that it was compensated by an increase in the supply of Antarctic-sourced deep water, at least in the New Zealand gateway for the Paciﬁc Deep Western Boundary Current (DWBC). Glacial conditions potentially favoured bottom water production through increased windiness plus an expansion of sea ice and associated brine rejection. Certainly Hall et al. (2001) record an increase in the speed of the Paciﬁc DWBC in glacial times as reﬂected by an increase in the size of ‘sortable silt’, a grain-size proxy for relative current strength (see also Venuti et al., 2007). For time frames longer than individual G-I cycles, the sortable silt data point to three phases of Paciﬁc DWBC vigour during the last 1.2 myr. These three phases cover the transition from a world dominated by 41 kyr climatic cycles to one where 100 kyr cycles prevailed and include: Phase 1 (1.2–0.87 Ma) of moderately strong ﬂow; Phase 2 or the MPT (0.87–0.40 Ma) of weaker ﬂow, and Phase 3 (0.40 Ma to present) when strongest ﬂows prevailed. The alternative hypothesis that little or no decline in NADW inﬂow occurred in glacial times is supported by Moy et al. (2006), Rosenthal et al. (1997) and others on the basis of various palaeoceanographic proxies including the distribution of d13C between ocean basins over G-I cycles. As NADW ﬂows to the Southern Ocean, its d13C content reduces due to remineralisation of organic carbon and mixing with southern-sourced waters with lower d13C. Depletion continues into the Indian and Paciﬁc oceans, the latter having the lowest d13C (Kroopnick, 1985). By comparing d13C proﬁles measured on benthic foraminifers from the main ocean basins bathed by NADW-sourced waters, it is possible to identify any variability of those waters. Moy et al. (2006) compared benthic d13C for cores spanning the last 180 kyr from the equatorial Atlantic, the Indian–Paciﬁc boundary at South Tasman Rise and the equatorial Paciﬁc. Equatorial Atlantic d13C was higher than at South Tasman Rise, but followed the same G-I cyclicity and maintained a gradient suggesting maintenance of northern-source deep water for the last 160 kyr. In addition, d13C data for the Tasman Rise and equatorial Paciﬁc were very similar highlighting a close tracking of Southern Ocean and central Paciﬁc waters. While the debate continues, most recent data (e.g. Moy et al., 2006; I. N. McCave, submitted and personal communication 2007) tend to favour a more subdued input of NADW and no major reorganisation of Southern Ocean water masses during glacial periods.

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11.4.1.2. Intermediate waters The dispersal of southern-sourced, intermediate depth waters is a key process in distributing heat, salt, nutrients and gases to equatorial and northern latitudes (e.g. Sarmiento et al., 2004). The northward transport of Subantarctic Mode Water (SAMW) and Antarctic Intermediate Water (AAIW) largely balances the southward export of NADW (Gordon, 1986). Both SAMW and AAIW form in the vicinity of the Subantarctic Front through buoyancy loss related to heat exchange, freshwater input and a northward Ekman transport although changes in wind stress and lateral advection may also play a role (e.g. McCartney, 1977; Rintoul et al., 2001; Toggweiler et al., 2006). While the mode(s) of intermediate water formation is not well understood, observations show that SAMW and underlying AAIW descend to B500 and B500–1,400 m depths, respectively. SAMW formation is fairly widespread, whereas AAIW generation is favoured in the SE Paciﬁc and SW Atlantic. Given their source at the wind-forced ocean surface, intermediate waters typically have elevated d13C contents as revealed by hydrographic data and intermediate depth benthic foraminifers, assuming other factors have not come into play (e.g. Lynch-Stieglitz et al., 1994; Pahnke and Zahn, 2005). In contrast, glacial benthic records show a marked reduction in d13C that is interpreted as a diminished supply of intermediate water. As outlined in more detail in the next section, the glacial Southern Ocean was (i) substantially cooler with sea-surface temperatures (SSTs) up to 61C colder than now (e.g. Mashiotta et al., 1999; EPICA Community Members, 2004; Barrows et al., 2007), (ii) more extensive following an equatorward migration of southern waters by B5–101 latitude (Howard and Prell, 1992; Gersonde et al., 2005), (iii) windier following an equatorward displacement and intensiﬁcation of zonal westerly winds (Shulmeister et al., 2004; Toggweiler et al., 2006) and (iv) supported a more extensive cover of sea ice B100% more than present (Gersonde et al., 2005). Pahnke and Zahn (2005) have argued that meltwater from the expanded sea ice, concomitant with its northward transport within an invigorated Ekman layer, reduced the surface density and the buoyancy-driven sinking of SAMW and AAIW. Yet despite this decline, intermediate waters remained a thermal conduit that linked Antarctica and the equatorial ocean judging by the strong coherence of their respective temperature records (Lamy et al., 2004; Kiefer et al., 2006). This linkage has been invoked by Weaver et al. (2003) as a potential control of meridional overturning in the North Atlantic. Pulses of meltwater from Antarctica could freshen AAIW, which upon arrival in the North Atlantic could potentially increase the density in contrast with surface waters to encourage sinking and the formation of NADW.

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11.4.1.3. Surface waters G-I cycles have a marked inﬂuence on the upper ocean by virtue of its direct interaction with the atmosphere. Reconstructions of SSTs for the last 25–400 kyr – the approximate range of records recovered by piston cores – highlight marked glacial cooling at orbital frequencies in the Atlantic sector of the Southern Ocean (e.g. Mortyn et al., 2002; Gersonde et al., 2003, 2005; Bianchi and Gersonde, 2004; Pahnke and Sachs, 2006), the Indian sector (e.g. Howard and Prell, 1992; Gersonde et al., 2005; Barrows et al., 2007), the SW Paciﬁc (Fig. 11.5; Weaver et al., 1998; Barrows et al., 2000, 2007; Pahnke et al., 2003; Neil et al., 2004; Pahnke and Zahn, 2005) and the SE Paciﬁc (e.g. Mashiotta et al., 1999; Lamy et al., 2004). A wide range of geochemical, isotopic and microfossil proxies is used to derive SSTs (see summary in Wefer et al., 1999), which can vary according to the proxy (Barrows et al., 2007). The quest for reliable comparisons between ocean basins is further confounded by some palaeoceanographic observations that record localised effects such as upwelling, rather than regional signals (e.g. Pahnke and Sachs, 2006). With such limitations in mind, the aforementioned studies reveal that ocean cooling was most pronounced around 40–461S during major glaciations (e.g. Marine Isotope Stages 2, 6, 8, 10, 12) when SSTs were 4–61C lower than present (Fig. 11.6). This change was recorded in most sectors of the Southern Ocean with the possible exception of the central Paciﬁc where a sparse database suggests that cooling was less severe (Gersonde et al., 2005). Certainly the western and eastern margins of the Paciﬁc Basin were subject to 4–61C coolings during the LGM (Lamy et al., 2004; Barrows et al., 2007). By comparison, the peak interglacials of Marine Isotope Stages 5e and 11 record SSTs up to 31CS warmer than present. The amplitude of SST swings reduces southwards (Howard and Prell, 1992; Gersonde et al., 2003) and once south of the Antarctic Polar Front, LGM SSTs, for example were B31C cooler than present at B501S and r11C cooler south of B521S, at least in the Atlantic-SE Indian sectors. SST records extending over a million years or more from the Southern Ocean are restricted mainly to DSDP and ODP sites, which because of their sparse coverage prevent identiﬁcation of regional patterns (Fig. 11.7). Nonetheless, they highlight temporal trends especially over the transition from 41 kyr-paced to 100 kyr-paced G-I cycles. Of note are SST records from (i) ODP 1090 located between the modern Subtropical and Subantarctic fronts in the SE Atlantic (Becquey and Gersonde, 2002; Hodell et al., 2002), and (ii) ODP 1123 positioned at the northern limit of the Southern Ocean, within the Subtropical Front in the SW Paciﬁc (Fig. 11.5; Crundwell et al., 2008). They were chosen because they are both based on foraminifer transfer

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Figure 11.6: (B) Stacked SST, normalised to a modern value of 01C, for the Southern Ocean, compiled by Barrows et al. (2007). It highlights the marked cooling of the LGM, when temperatures were up to 51C cooler than now and the abrupt swings in SST that are near-synchronous with warm phases in Antarctica (C) but mainly leading prominent D-O events in the Northern Hemisphere (A). functions and come from open-ocean sites beyond local complexities imposed by bathymetry. Both records exhibit the same three phases of SST variability (Fig. 11.7), but the type and degree of variability at times differ in response to the regional oceanography. Phase 1: 41 kyr cycles (1.2 Ma and earlier to 0.87 Ma). SW Paciﬁc temperatures exhibit high amplitude ﬂuctuations of 4–71C within a G-I range of 9–181C. These ﬂuctuations are superimposed on an overall cooling trend that is especially evident with interglacial optima, which decline progressively from 18 to 141C. Contemporaneous SE Atlantic SSTs have more restricted amplitudes of 1–31C and G-I range of 3–51C. A cooling trend is less clear,

Figure 11.7: Composite plot of Antarctic and Southern Ocean climate proxies spanning the last 1.2 myr. The data allow the coupled G-I variability within the atmospheric, oceanic and cryospheric systems to be evaluated in the context of longer-term changes extending from the 41 kyr climate system (Phase 1) through the MPT (Phase 2) and into the 100 kyr climate system (Phase 3).

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but an irregular decline is evident from B1.06 to 0.87 Ma when interglacial optima went from 7 to 41C. The general cooling at both sites reﬂects an expansion of the Antarctic cryosphere accompanied by a northward migration of the cold surface waters, while at depth there may have been a general increase in the thermohaline circulation at least into the Paciﬁc Ocean (Hall et al., 2001). Thus, at ODP 1123, Subtropical Water was displaced by Subantarctic Water with the former returning at the last part of the terminations and early interglacials. In contrast, the more southerly sited ODP 1090 was mainly affected by circumpolar surface waters through much of the G-I cycles thus accounting for the colder temperatures with subdued amplitudes and G-I ranges. Phase 1 terminated at both sites at the prominent interglacial of Marine Isotope Stage 21. Phase 2: MPT (B0.87–0.4 Ma). This fundamental transition from a 41–100 kyr dominant world was marked at both sites by persistently warmer and more prolonged interglacial periods that were interrupted by some of the coldest glacial periods in the Pleistocene. G-I cyclicity was most marked in the SE Atlantic, which came under the inﬂuence of Subantarctic Water and even Subtropical Water during Marine Isotope Stage 15. Not surprisingly, that southward migration of Subtropical Water was particularly well shown at ODP 1123 where it dominated SSTs from Marine Isotope Stages 15–13 with the intervening glacial, Marine Isotope Stage 14, bringing only a modest cooling (o21C). The general warming and southward migration of warmer waters during the MPT is consistent with a reduced Antarctic inﬂuence, which is supported by a less vigorous thermohaline circulation into the Paciﬁc as suggested by Hall et al. (2001). The marked G-I variability is interpreted by Becquey and Gersonde (2002) to reﬂect the increasing inﬂuence of 100 kyr cycles, which in turn are driven primarily by changes in NH ice sheets. However, the impact of these drivers appears to be blunted by subtropical inﬂuences as attested by ODP 1123 record especially over Marine Isotope Stages 15–13. Phase 3: 100 kyr cycles (B0.4 Ma to present). In the SE Atlantic the onset of the 100 kyr world was accompanied by highly variable SSTs with G-I contrasts of up to 81C. These major ﬂuctuations occurred against a background of mainly warmer temperatures at ODP 1090, which lay mainly within Subantarctic Water. While also exhibiting 100 kyr-paced SSTs, the G-I differences at ODP 1123 are less marked with all but one (Marine Isotope Stages 6-5) less than 51C. Such a muted response may reﬂect the inﬂuence of the south Paciﬁc Subtropical Inﬂow (e.g. Carter et al., 2008). Despite differences in SST amplitudes, the SE Atlantic and SW Paciﬁc data reveal a progressive reduction of B31C in interglacial optima over Phase 3.

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11.4.1.4. Ocean fronts From the preceding discussion it is clear that the variability of SSTs over G-I cycles is intimately linked to the meridional displacement of polar waters, which is affected by the state of the Antarctic cryosphere and zonal winds. Glaciations witnessed a major expansion of winter sea ice. In the LGM, sea ice grew by B100% compared to now, extending 3–81 latitude in the Atlantic, 7–101 in the Indian and possibly B2–51 in the Paciﬁc, bearing in mind the last region has sparse data coverage (Gersonde et al., 2005). This expansion was accompanied by an intensiﬁcation and equatorward displacement of zonal winds (Thiede, 1979; Shulmeister et al., 2004) along with a northward migration of cold water and ocean fronts (e.g. Howard and Prell, 1992; Becquey and Gersonde, 2002; Gersonde et al., 2003; Crundwell et al., 2008). The extent of the migrations depend upon the strength and duration of the climatic drivers, as well as the morphology of the Southern Ocean ﬂoor, for example any G-I displacement of fronts off eastern New Zealand, is limited by extensive submarine elevations such as the Chatham Rise where the STF is restricted to 1–21 latitude along the Rise crest (Weaver et al., 1998; Sikes et al., 2002). In contrast, frontal systems and associated water masses in the open Southern Ocean are free to migrate further aﬁeld (Fig. 11.5). However, it is an open question as to how SSTs represent shifts in ocean fronts. As pointed out by Gersonde et al. (2005), control sites are too scattered to delineate the temperature and salinity gradients that deﬁne frontal systems. Accordingly, the following outline of frontal migrations is indicative only. The data of Becquey and Gersonde (2002), Gersonde et al. (2005) and Howard and Prell (1992) indicate northward migrations of 41, 5– 101 and 2–31 for Antarctic Polar Front in the Atlantic, Indian and Paciﬁc sectors, respectively, during the LGM. At the same time, the SAF shifted 4–51 and 5–101 in the Atlantic and Indian, respectively. SSTs from the easternmost (Lamy et al., 2004) and westernmost Paciﬁc (Crundwell et al., 2008) suggest similar northward displacements. In contrast, migrations of the STF were restricted to only ca. 2–31 and 51 latitude in the Atlantic and Indian sectors suggesting an intensiﬁcation of oceanographic gradients at the northern limit of the Southern Ocean (Gersonde et al., 2005). The reasons for such limited migration are not clear but it maybe reﬂect a counteraction by the subtropical gyres in the major basins. In addition to the G-I cycles, long-term records from the ODP sites (Fig. 11.7) identify broad changes in the fronts. At ODP 1090, the PF prevailed, representing a possible shift of B71 latitude north of its present position (Becquey and Gersonde, 2002). During the MPT the site was dominated by the SAF, but with interruptions by glacial excursions of the

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PF. In Phase 3, the SAF again dominated but was interspersed by interglacial incursions of the STF. In the SW Paciﬁc, the more northerly located ODP 1123 spent most of the last 1.2 myr within or north of the STF (Crundwell et al., 2008) except during peak glacial periods. Taking the mean annual SST of the northward edge of the modern SAF to be 111C (Uddstrom and Oien, 1999), then the mean annual palaeo-SST proﬁles reveal that (i) Phase 1 was a time of frequent glacial incursions of the SAF and associated waters representing a northward migration of B51 latitude; (ii) Phase 2 MPT witnessed only infrequent SAF migrations, which ceased in the latter half when the STF prevailed and (iii) the SAF was located well south reaching Site 1123 only in the major glaciation of Marine Isotope Stage 6.

11.4.1.5. Surface ocean currents As noted in the previous section, SST data indicate northward migrations of the PF and SAF during glacial periods. As these fronts accommodate most of the ﬂow within the ACC (e.g. Whitworth, 1988; Rintoul et al., 2001) the inference is that this major current system also moved north, possibly by B51 or more of latitude. However, in light of (i) limitations of SST data to deﬁne past frontal positions, (ii) the marked topographic steering of the modern ACC by mid-ocean ridges (e.g. Gordon et al., 1978; Orsi et al., 1995; Moore et al., 1999) and (iii) the strong latitudinal variability of modern ACC fronts such as the PF, which ranges over 5–71 of latitude (Moore et al., 1999), it is unclear how the ACC behaved. Uncertainty also arises as to whether the ACC strengthened under the stronger wind regimes of glacial periods. Certainly the ACC is mainly wind-driven, with the extent of this forcing dependent upon the position of mid-latitude westerly winds (Toggweiler et al., 2006). Wind strengthens the ACC by direct shear but also enhances the northward Ekman transport, which in turn is compensated by the southward transport and eventual upwelling of deep water at the Antarctic margin. These effects are most pronounced when winds are aligned with the ACC as is the case today. Although glacial period westerly winds (Moreno et al., 1999; Sigman and Boyle, 2000) and the main ACC fronts were located further north than today, it is not known if winds and current were aligned. However, where migration of the ACC is restricted by the seabed morphology as in the Scotia Sea (Pudsey and Howe, 1998), or off eastern New Zealand (Neil et al., 2004), the ACC appears to have strengthened, although it is unclear if this was a response to increased wind-forcing, compression of fronts against the bathymetry or both these processes.

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11.4.1.6. Iceberg discharge and migration A proxy for G-I behaviour of Antarctic ice is found in the IRD records of Southern Ocean sediments. ODP Site 1011, off the Antarctic Peninsula provides a generalised IRD history spanning 3 Ma (Cowan, 2002). Features are phases of pronounced IRD deposition around 2.8, 1.9 and 0.85 Ma, interleaved by more subdued deposition that followed G-I cycles with a frequency of 41 kyr from 2.2 to 1.0 Ma and 100 kyr frequency after 0.4 Ma. In the latter cycles, the late glacials and following interglacials were the main times of ice rafting. A more detailed record for the last 300 kyr in the Weddell Sea conﬁrms the strong IRD signal at G-I transitions, but shows high deposition of IRD continuing into the early part of the following glacial period (Fig. 11.8; Grobe and Mackensen, 1992). The Weddell Sea record also reveals a consistent decline in IRD production since 200 kyr

Figure 11.8: Normalised stacked proﬁles for IRD generated at source in the Weddell Sea (Grobe and Mackensen, 1992) and deposited at a distant depocentre – the Campbell Plateau, New Zealand (Carter et al., 2002). Both proﬁles show similar broad outlines but the key difference is that the times of maximum IRD production are not faithfully reproduced off New Zealand especially during a G-I transition when a warming ocean and reducing windiness would be less favourable for iceberg preservation.

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(Marine Isotope Stage 7). Superimposed on the G-I cycles are IRD pulses of millennial-scale frequency which, given uncertainties of age models, are traceable from the SE Atlantic to SW Paciﬁc, at least for the past 70 kyr (Labeyrie et al., 1986; Kanfoush et al., 2000; Carter et al., 2002). The Kanfoush et al. (2000) IRD events coincide with periods of NADW production and warm interstadials in the North Atlantic. This led to the proposition that the increased inﬂow of relatively warm NADW and/or a rise in sea level destabilised ice sheets to promote iceberg discharge. Once calved, icebergs from the Weddell to Ross Sea regions were most likely captured within the Weddell, Ross and un-named coastal gyres judging by modern iceberg trajectories (Keys, 1990). Nevertheless, some were eventually entrained by the ACC to move east as identiﬁed from modern iceberg paths (Tchernia and Jeannin, 1983), last glacial IRD dispersal patterns (Cooke and Hays, 1982), and a meltwater signal generated about 35–17 kyr (Labeyrie et al., 1986). Not surprisingly, the Weddell Sea IRD proﬁles (Grobe and Mackensen, 1992) broadly correlate with those off eastern New Zealand (Fig. 11.8; Carter et al., 2002). However, there are also differences in that the interglacial records are more subdued off New Zealand, a feature Carter et al. (2002) attributed to iceberg melting en route from Antarctica. In contrast, glacial records are better correlated indicating better preservation of travelling icebergs on account of colder glacial seas and possibly an intensiﬁed ACC (Pudsey and Howe, 1998; Neil et al., 2004). 11.4.2. Millennial-Scale Cycles Antarctic ice cores record frequent abrupt changes in temperature at millennial scales that are difﬁcult to explain by variations in orbital parameters. Over the last 90 kyr, Antarctica has been subject to seven prominent warm periods designated A1–A7 (Blunier and Brook, 2001, and discussed above). With the possible exception of A5 and A6, the others are well expressed in SST records from southern mid-latitudes (Fig. 11.6; Barrows et al., 2007). But perhaps the best-documented abrupt change in terms of identifying oceanic responses is the Antarctic Cold Reversal (ACR), which is timed at around B14.1–12.4 ka according to the EPICA deuterium record (Blunier et al., 1997; EPICA Community Members, 2004). An abrupt cooling of B21C appears to have been accompanied by an expansion of sea ice (Shemesh et al., 2002) and a modest intensiﬁcation of winds as identiﬁed from sea salt and dust proxies in ice cores (Stenni et al., 2001; Ro¨thlisberger et al., 2002). As noted previously, an expanded Antarctic cryosphere can invigorate the THC (e.g. Hall et al., 2001), and judging by the modern oceanography,

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such a change would be translated rapidly through the Southern Ocean (e.g. Haine et al., 1998). Thus, off eastern New Zealand, the Paciﬁc gateway for the THC, benthic d18O becomes heavier directly in phase with the ACR (Carter et al., 2008). In contrast, change in the surface ocean about the time of the ACR is variable. South of the Subtropical Front, SSTs cool in phase with the ACR, as recorded in the SW Paciﬁc (Pahnke et al., 2003), SE Paciﬁc (Lamy et al., 2004) and the south Atlantic (Kanfoush et al., 2000; Sachs et al., 2001; Shemesh et al., 2002; Gersonde et al., 2003). These sites have direct atmospheric and oceanic links to Antarctica, especially during glacial periods when polar effects are enhanced through the northward migration of westerly winds and surface ocean waters. By comparison, mid-latitude records have a more delayed and muted response to the ACR. Off eastern New Zealand, for example, there was no obvious response until B13.5 ka when the ACR was at its coldest (Carter et al., 2008). Then, SSTs became cooler, marine fertility dropped and the uppermost ocean either became more mixed or the thermocline more shallow. Onshore, pollen and speleothem data reveal cooler, windier conditions that were accompanied by an expansion of glaciers and, at the coast, by a likely stillstand in sea level (Carter et al., 1986; Turney et al., 2003; Williams et al., 2004). The delayed reaction to the ACR is likely connected to a contemporaneous re-establishment of the Subtropical Inﬂow as it migrated south during the deglacial phase (e.g. Martinez, 1994). Like the southwest Paciﬁc, the south Indian Ocean also cooled out of phase. SSTs reduced by 0.81C between 13.2 and 12 ka, B1,000 years after commencement of the ACR demonstrating the regionality of abrupt climate change effects (Stenni et al., 2001). As core chronologies improve (see Steig, 2001), it is becoming apparent that some millennial-scale variability in the Southern Ocean is linked with abrupt changes documented in the NH. Barrows et al. (2007) draw attention to the timing of Southern Ocean warm phases, A1–A4, with warm phases of D-O cycles, 8, 12, 14 and 17 identiﬁed in the Greenland ice core, GISP 2. Likewise, Sachs and Anderson (2005) observe phases of marked algal productivity off eastern New Zealand that occurred within 1–2 ka of massive iceberg inﬂuxes associated with Heinrich Events 1–6 in the North Atlantic. Another northern cold event that has received much attention in the Southern Ocean is the Younger Dryas (YD) of 13–11.5 kyr (e.g. Morigi et al., 2003; Turney et al., 2003; Bianchi and Gersonde, 2004; Sachs and Anderson, 2005; amongst others). Because of the YD’s partial overlap with the ACR (14.1–12.4 kyr) and the Oceanic Cold Reversal (13.2–12 kyr) in the Indian Ocean (Stenni et al., 2001) conclusive evidence for a YD effect is still under debate, the resolution of which hinges on the quality of palaeoenvironmental chronologies. Once resolved, the mechanisms that translate YD and other

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abrupt NH perturbations to the Southern Ocean (and vice versa) can be better evaluated to allow a more informed assessment of potential change in the future.

11.5. Modelling of Pleistocene Ice Volume Variations 11.5.1. Northern Hemisphere, the Last Million Years Variations in global ice volume during the Pleistocene, and probably the Pliocene, have been dominated by growth and decay of Northern Hemispheric ice, and most ice-sheet modelling of this period has been concerned with the Laurentide, Cordilleran, Greenland and Eurasian ice sheets. Weertman (1976) was the ﬁrst to include ice dynamics in model simulations of long-term ice-sheet evolution. He used a perfectly plastic icesheet proﬁle representing a north–south cross section of the Laurentide Ice Sheet, and forced it from 400 ka to the present with a prescribed pattern of net annual accumulation minus ablation versus latitude and elevation (termed a snowline pattern below), shifted vertically in proportion to summer insolation at 651N to represent actual Milankovitch orbital variations. This study was the ﬁrst to capture the several thousand-year lag of ice volume behind climate forcing due to the mass inertia of the ice sheet. It produced reasonable B23–41 kyr cycles in direct response to precessional and obliquity orbital variations, but did not realistically simulate the dominant observed 100 kyr ice cycles. Subsequent work extended this study to (i) use energy-balance climate models (EBMs) instead of a prescribed snowline pattern (Pollard, 1978; Berger et al., 1990; Deblonde and Peltier, 1993) and (ii) use ice-sheet ﬂow models instead of plastic proﬁles, allowing bedrock topography and lagged depression below the ice (Birchﬁeld et al., 1981; Pollard, 1982; Hyde and Peltier, 1987). For the most part, the addition of climate models did not change Weertmans’ results signiﬁcantly, and just showed that the smooth summer-temperature variations in EBMs have much the same effect as a shifted snowline pattern. The additions of high-latitude topography and lagged bedrock depression and rebound were found to be more important: high-latitude plateaus allow nascent ice caps to initiate more easily following interglacials, and deep bedrock depressions at glacial maxima amplify subsequent retreat due to warmer air temperatures at lower elevations. Although these mechanisms produced more non-linear response, they were still not enough to produce realistic 100 kyr response and full G-I cycles.

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Retreats are triggered in these models by hot northern summer orbital conﬁgurations that are coeval with the start of major observed terminations, but the modelled retreats are not complete, allowing substantial ice cover to persist through interglacials. It was found that some additional physics were needed that come into play during the major deglaciations to terminate each 100 kyr cycle, and a large variety of mechanisms have been proposed. These range from cyclic freezing versus melting and surging at the ice-sheet base (Oerlemans, 1982b; MacAyeal, 1993), calving by proglacial lakes (Pollard, 1982), prescribed climatic effects of atmospheric CO2 and/or North Atlantic heat transport and thermohaline circulation (Tarasov and Peltier, 1997), and atmospheric dust concentrations (Peltier and Marshall, 1995). These mechanisms involve or respond to some other long-term component (besides ice volume) with intrinsic time scales of thousands to tens of thousands of years, which is logically necessary because mechanisms that only involve fast-time-scale components (atmosphere, snow, sea ice, upper ocean) would have been triggered not just during major deglaciations but at other times with similar orbits and mid-range ice extents. However, there are two exceptions: snow aging (Gallee et al., 1992) and sea-ice switching (Gildor and Tziperman, 2001). Other aspects of the 100 kyr cycles have been studied recently, such as phase locking by eccentricity or obliquity/precession (e.g. Tziperman et al., 2006) and whether these cycles are deterministic at all (Wunsch, 2003). Although many of the models mentioned above produce realistic Pleistocene ice volume time series including B23, 41, and 100 kyr cycles, there is still no consensus on what processes are responsible for the rapid inceptions and complete deglaciations that are key to the dominant 100 kyr cycles of the last million years. An entirely different approach is to use zero-dimensional ‘box’ models comprised of a few coupled non-linear ordinary differential equations for quantities such as global ice volume, mean ocean temperature and atmospheric CO2 level (Matteucci, 1989; Saltzman and Verbitsky, 1993; Paillard, 1998). These can be tuned to yield fairly realistic ‘ice volume’ time series forced by Pleistocene orbital variations. Hargreaves and Annan (2002) used Saltzman’s model and Monte Carlo Markov Chain (MCMC) techniques to ﬁnd the optimal parameter set and probability density functions yielding the most realistic results; similar approaches have used three-dimensional ice sheet models (Tarasov and Peltier, 2004). Paillard’s zero-dimensional results basically support the general conclusion from more physical models that nonlinear thresholds are needed to successfully produce 100 kyr cycles from the higher-frequency orbital forcing. However, the highly conceptual formulation of the zero-dimensional models make it hard to translate their results into speciﬁc action items for improvements in more physically explicit models.

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A unique approach was taken by Bintanja et al. (2005) to address the ambiguity between ice volume and local sea-water temperature inherent in deep-sea core d18O records. They used a NH ice-sheet ﬂowline model, driven through multiple G-I cycles over the last million years with simpliﬁed climate–orbital linkages as in many previous studies. But instead of simply running the model forward in time, (i) a standard empirical formula was used to predict foraminiferal d18O at any time from the model’s current deep-sea temperature and ice volume, and (ii) at each time step the climate temperature was determined not from a physical model, but from the requirement that, when the ice model was stepped forward, the computed foraminiferal d18O yielded the observed deep-sea-core value for the next time step. This produced a self-consistent estimate of the relative contributions of Northern Hemispheric ice versus deep-ocean temperature in d18O core records, with the main conclusion that the ice-sheet contribution varied from B10% in the beginning of glacial cycles to B60% at glacial maxima; the latter is consistent with relict pore water measurements (Schragg et al., 1996). Some attempts have been made to use General Circulation Models (GCMs) in conjunction with long-term ice-sheet variations. Computer time is prohibitive for direct GCM integrations longer than a thousand years or so, and several strategies have evolved to circumvent this limitation. Several studies have driven three-dimensional ice sheet models through the last glacial cycle (last B125 kyr), with climates obtained by interpolating between two stored GCM snapshots at the LGM (21 ka) and modern. The weighting of the two GCM climates is proportional to an observed ice core record (usually Greenland GRIP) representing NH climate of the last 125 kyr (Marshall and Clarke, 1999b; Zweck and Huybrechts, 2005; Charbit et al., 2007). However, this procedure heavily constrains the predicted variations of ice volume (on millennial, orbital and 105-year time scales) to those in the ice core record, so the approach is useful mainly for diagnostic assessment of processes in the ice model. For instance, one striking diagnostic result is the greatly expanded extent of basal melting under the Laurentide as LGM is approached, which is consistent with the basal freeze-thaw-surge hypothesis for ampliﬁcation of 100 kyr cycles (Marshall and Clark, 2002; cf. Johnson and Fastook, 2002). To couple relatively coarse-grid GCM climates (B100’s km) to the ﬁner ice-sheet model grids and topographies (B10’s km) in these experiments, straightforward downscaling techniques have been developed (e.g. Ramstein et al., 1997; Thompson and Pollard, 1997; Marshall and Clarke, 1999a). First, monthly mean air temperatures and precipitation are interpolated both horizontally and vertically to the ice surface locations. Degree–day parameterisations are used to convert monthly temperatures to snow or ice melt, which are able to capture subtle effects of orbital changes in seasonal

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insolation (e.g. Huybers, 2006). Some calculations explicitly account for refreezing of meltwater. Also, ‘anomaly’ techniques are often used to reduce biases in GCM climate, by only using GCM predicted changes in climate from the present, combined with modern observed climatology. This is usually done additively for temperature, and by ratio for precipitation to avoid negative precipitation values. A few studies have used asynchronous coupling between GCMs and ice sheets, whereby the GCM is run for a few decades every few thousand years and its climate is used to drive the ice sheet model alone for the next few thousand years. Then the new ice-sheet extent is used to update the GCM’s surface conditions, and the sequence is repeated. For the Pleistocene, this has been done only over the last 21 kyr through the last deglaciation (Charbit et al., 2002). More recently, Earth Models of Intermediate Complexity (EMICs) have made it possible to run coupled climate–ice sheet models through long periods (Charbit et al., 2005). Realistic Pleistocene cycles have been obtained in this way (Calov and Ganopolski, 2007), but, as found in some of the earlier studies above, prescribed variations of atmospheric CO2 and North Atlantic ocean heat transport were needed to obtain complete 100 kyr cycles. A new approach using climate parameterisations based on past GCM snapshots has recently been taken by Abe-Ouchi et al. (2007). Many Global Climate Model (GCMs) and Regional Climate Model (RCMs) have performed ‘snapshot’ simulations of Pleistocene climates at individual times, usually with prescribed ice sheets. One of the ﬁrst such applications was Gates (1976) at the LGM, using prescribed CLIMAP SSTs and ice sheet reconstructions; recently, fully coupled A/OGCMs have been used for LGM (Hewitt et al., 2003; Kim et al., 2003), and GCMs and EMICs have been applied to other times such as the end of the last interglacial, sometimes driving nascent ice sheets (deNoblet et al., 1996; Vettoretti and Peltier, 2004; Kubatzki et al., 2006). A few studies have used RCMs or zoomed-region GCMs to simulate the mass balance over NH ice sheets at LGM and other times (Hostetler et al., 2000; Krinner et al., 2004). A detailed survey of these simulations is outside the scope of this chapter, but two general points emerge concerning ice sheets: (i) there is wide scatter between GCMderived mass balances of ice sheets, due to the strong sensitivity of summer melt to air temperatures over the ablation zones (Pollard and PMIP, 2000), and (ii) although reasonable mass balance patterns on major NH ice sheets can be attained, several characteristics consistently disagree with geologic data, such as too much ice buildup over Alaska and Siberia (Charbit et al., 2007). It is relatively easy to adjust the downscaling of climate or other parameters to achieve a reasonable mass balance for a given GCM, time and ice sheet, but

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much more challenging to achieve good results for other times and other ice sheets using exactly the same model and parameter values. 11.5.2. Mid-Pleistocene Transition Before the MPT at around 900 ka BP, the 100 kyr cycle was essentially absent in d18O deep-sea core records, and most of the spectral power was at the orbital obliquity period of B41 kyr, with little at the precessional period B23 kyr (Mudelsee and Schulz, 1997). A number of models that explicitly involve NH ice sheets have attempted to understand this abrupt change. Most rely on a special mechanism that allows 100 kyr cycles in the model only when the maximum ice sheet size is large, and not before B1 Ma when the ice sheets never reached the required size. This size evolution is accomplished in turn by an assumed long-term trend in some external model parameter such as declining atmospheric CO2 (Deblonde and Peltier, 1991; Mudelsee and Schulz, 1997; Berger et al., 1999). Clark and Pollard (1998) achieved the same result not by a prescribed long-term trend, but by adding a deforming sediment component underneath the ice sheet. Starting with a 50 m regolith at 3 Ma, the early ice sheets were kept thin by the greater basal sliding due to the sediment. After the sediment is eroded by repeated glacial growth and retreat, by about 1 Ma enough hard bedrock is exposed to support the greater basal shear stresses of thick ice sheets, allowing the 100 kyr mechanism in that model (proglacial calving) to operate. However, the true cause(s) of the MPT remain unknown. Other proposed mechanisms are noted below in section 11.6.4. None of these models address the dominance of B41 kyr spectral power (driven by obliquity) compared to B23 kyr (driven by precession) before the MPT. One recent theory (Raymo et al., 2006) hypothesises that before the MPT, the EAIS was somewhat smaller than at present and much more variable, with terrestrial margins and ablation (summer-melting) zones that responded signiﬁcantly to orbitally driven summer-temperature variations. If so, precession-driven ice volume variations would have been out of phase between the hemispheres, and obliquity-driven variations in phase, resulting in a cancelling between northern and southern ice volume variations at 23 kyr, and dominant 41 kyr power. This hypothesis is testable by establishing the character and variability of the Antarctic Ice Sheet during the Pliocene up to the MPT (see below). Another recent hypothesis notes that total melt season insolation is affected more strongly by obliquity than by precession (Huybers, 2006); this is testable by models with seasonal cycles and positive-degree-day parametrisations of ice melt.

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11.5.3. Millennial-Scale Variability Most climate and ice sheet modelling that addresses Pleistocene millennialscale variability has been in the form of snapshots for individual times or modes, and not in long-term time sequences. This is true for mechanisms involving see-saw connections, ocean response to meltwater discharge, and thermohaline interactions between southern and northern high latitudes (reviewed in Rahmstorf, 2002; Clark et al., 2007). One exception is the modelling of the ice-sheet binge-purge mechanism for Heinrich Events (HEs) by Calov et al. (2002) who used a three-dimensional Laurentide ice sheet coupled to an EMIC. They integrated the coupled system through 200,000 years and obtained quasi-periodic HE-like surges due to basal melting and refreezing over Hudson Bay and Strait. However, the true causes of HEs and whether they are internal ice sheet oscillations or manifestations of oceanic sea-saw behaviour are active research questions (Clark et al., 2007).

11.5.4. Antarctica Little data are available that directly records Antarctic Ice Sheet variations during the last few million years. Most climate and ice sheet modelling of Pleistocene Antarctica has been directed at individual times: LGM, modern and future (e.g. Oerlemans, 1982a; Warner and Budd, 1998; Huybrechts et al., 2004). Throughout the Pliocene and Pleistocene the EAIS is usually thought to have been stable with few deeply submerged marine margins, narrow continental shelf areas limiting expansion (Anderson et al., 2002) and marginal summer temperatures that are too cold to allow signiﬁcant surface melt even with favourable orbital conﬁgurations. One regional exception is the Prydz Bay–Lambert Graben drainage system, where geologic and seismic studies have found signiﬁcant regional variations in glacial extent through the later Cenozoic (McKelvey et al., 2001; Whitehead et al., 2004; O’Brien et al., 2007). On continental scales, though, the EAIS has been insensitive to the B120 m sea-level variations caused by Northern Hemispheric ice cycles, and by orbitally and CO2-induced local air temperature changes. Contrary views have been proposed by Raymo et al. (2006) involving smaller and more variable EAISs prior to the MPT (B1 Ma, see above), and by several researchers (e.g. Webb and Harwood, 1991) who suggest that Sirius Group deposits in the Transantarctic mountains imply drastic EAIS retreat in the Wilkes Bay sector during the Pliocene (see Miller and Mabin, 1998).

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In contrast, the WAIS is grounded well below sea level with extensive deep grounding lines and wide continental shelves in the Weddell and Ross embayments. Consequently it is much more sensitive than the EAIS to variations in sea level and ocean temperatures via ice shelf basal melt. The WAIS is thought to have repeatedly expanded to the continental shelf limits and back during the Pleistocene, driven primarily by sea-level variations of B120 m caused by NH ice growth and decay. A few three-dimensional Antarctic ice models have been run in this way through the last 400 kyr (Ritz et al., 2001; Huybrechts, 2002), forced by prescribed temperature and sea-level variations from ice core and deep sea core records. These threedimensional ice models have the capacity to some degree to simulate ﬂoating ice and advancing and retreating marine grounding lines. They capture what is known of the WAIS and EAIS behaviour since the LGM reasonably well, including the retreat of WAIS grounding lines over the Ross and Weddell embayments (Conway et al., 1999; Anderson et al., 2002), as illustrated in Fig. 11.9. However, the models’ treatments of basal hydrology and grounding line migration are simpliﬁed and open to question, especially in the light of recent theoretical developments in grounding line treatment that suggests coarse-grid models are not capable of realistic migration (Schoof, 2007), and fuller-stress models are needed to capture drawdown in ice stream regions (Payne et al., 2004). Most of the variation in total grounded ice volume in the Huybrechts and Ritz glacial-cycle models is due to expansion and contraction of grounding lines across continental shelves, mostly in the WAIS Ross and Weddell sectors, but also to the west of the Antarctic Peninsula and other areas. The modelled G-I (BLGM vs. modern) change in Antarctic ice volume is equivalent to B15–20 m of sea level. The EAIS interior responds in the opposite sense, contracting slightly at Northern Hemispheric glacial maxima due to lower model snowfall rates (which are reduced as temperatures fall). At these times, modelled surface elevations in the central EAIS plateau decrease by B100–150 m compared to the present, with minor volume changes equivalent to just a few metres of sea level. Just as for the NH, more sophisticated regional climate, ocean and ice models have been applied to Antarctica but only for particular ‘snapshot’ times, mainly the present and LGM. These include RCMs (Bailey and Lynch, 2000; Bromwich et al., 2004), ice-sheet models including basal sediment and hydrology (Vogel et al., 2003), longitudinal stresses (Payne et al., 2004), subglacial lakes (Siegert, 2005; Pattyn and Siegert, 2007) and regional ocean models capable of resolving sub-ice shelf circulation (Jenkins and Holland, 2002; Holland et al., 2003). The combination of additional physical processes and higher model resolution could well be important for

Figure 11.9: Snapshots of Antarctic Ice Sheet Evolution during the last G-I cycle, simulated by a three-dimensional ice-sheet model driven by parameterised forcing, from Huybrechts (2002; his Fig. 3). Shown is surface ice elevation relative to present sea level. Contour interval is 250 m; thick lines are for every 1,000 m; the lowest contour is for 250 m and generally close to the grounding line (s.l.e., sea-level equivalent). The x- and y-axis-coordinates are arbitrary spatial values used by the model.

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the long-term evolution of Antarctic ice, but their use in time-evolving simulations awaits future work.

11.6. Synthesis: Antarctic Climate Evolution Since B3 Ma 11.6.1. Late Pliocene Cooling, Ice Sheet Expansion and the Development of a Marine Ice Sheet Margin Around Antarctica In the Weddell Sea, Kennett and Barker (1990) were the ﬁrst to identify a cooling step marked by a reduction in sedimentation rate, lower diatom abundance and the presence of sea-ice diatoms during the Late Pliocene. Along the Paciﬁc margin of the Antarctic Peninsula, seismic data correlated into ODP Leg 178 drill cores, documented a change in depositional style associated with reduced sediment supply and meltwater volume as the Antarctic Ice Sheets expanded and began cooling towards their present state from around 3–2.5 Ma (Rebesco et al., 2006). A large spike in IRD abundance at 2.8 Ma recorded in Antarctic Peninsula continental rise Site 1101 core also supports the development of a marine ice terminus along the edge of the WAIS (Cowan, 2002). In the Ross Sea, Bart (2004) interpreted a major glacial unconformity in seismic reﬂection data of inferred Late Pliocene age as evidence for widespread westward expansion of the WAIS ice streams into the Ross Sea. Although some uncertainties surround the chronology of DVDP cores in the mouth of Taylor Valley, and CIROS-2 cores in Ferrar Fiord, expansion of WAIS-sourced ice into the Dry Valleys also appears to have occurred about this time (McKelvey, 1981; Powell, 1981). Based on the above and new evidence from the ANDRILL AND-1B record (Naish et al., submitted) it appears that between 3.0 and 2.5 Ma, highlatitude climate cooling drove both the WAIS and EAIS towards their present expanded cold polar state. Relatively warm and often terrestrially based ice margins were replaced by more permanent marine termini and the development of ice shelves.

11.6.2. Late Pliocene–Early Pleistocene G-I Climate Cycles and the Role of Northern Hemisphere Glacio-Eustasy Direct physical sedimentary records of Antarctic Ice Sheet variability (e.g. IRD, proximal glacimarine cycles) and more distal ocean records of sea-ice

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distribution (e.g. diatom palaeoecology), thermohaline circulation (e.g. grain size), ocean temperatures (e.g. d18O), frontal dynamics and surface circulation (e.g. faunally derived assemblages and SSTs) all show a strong covariance with the 41 kyr cycle in Earth’s obliquity prior to the last 800 kyr. It has been proposed that long-term variations in the duration of summer insolation controlled by orbital obliquity (Huybers, 2006), generally dominates over the inﬂuence of the intensity of NH insolation (e.g. Raymo et al., 2006) prior to the Mid-Pleistocene. Marine Isotope Stage 31 may be the exception (Scherer et al., 2008). Strong coupling between the different elements of the SH ice-sheet–ocean-climate system, supports a primary obliquity inﬂuence on G-I variability between 3 and 1 Ma. Glacial periods result in the northward expansion of seasonal sea ice, SSTs up to 61C colder than now, equatorward migration of frontal zones by B5–101 latitude (Howard and Prell, 1992; Gersonde et al., 2005), equatorward displacement and intensiﬁcation of zonal westerly winds (Shulmeister et al., 2004; Toggweiler et al., 2006), invigorated surface circulation (ACC) and intensiﬁed abyssal currents (e.g. Hall et al., 2001). All of these processes appear to have occurred 3–7 kyr before the d18O ice volume maximum (e.g. Crundwell et al., 2008) when Antarctic Ice Sheets are fully extended onto the continental shelf. The converse is true for interglacials. Stratigraphic analysis of 41 kyr duration, Late Pliocene glacimarine cycles in the AND-1B core straddling the Gauss–Matuyama (G-M) polarity transition (ca. 2.6 Ma) display abrupt transitions between ice-proximal diamictites and open marine biosiliceous deposits, implying rapid oscillations between glacial and marine environments with signiﬁcant volume changes in the WAIS (Naish et al., 2008). Intriguingly, the facies also imply a cooler style of ice sheet compared with the Early Pliocene cycles, indicating reduced amounts of subglacial meltwater and terrigenous sediment (McKay et al., accepted). From B2.6 Ma polar Antarctic ice volume changes may have been controlled by the effect of NH glacio-eustasy on its marine margin, and that this mechanism accounts for much of the orbital variability since 2.6 Ma (cf. Raymo et al., 2006).

11.6.3. Marine Isotope Stage 31: Anomalous Continental-Scale Warmth An enigmatic interval of foraminiferal ooze and coccolith-bearing assemblages in the Weddell Sea and Prydz Bay cores together with a bioclastic limestone in the Ross Sea at B1 Ma, imply a signiﬁcant warming and change in ocean chemistry around the periphery of Antarctica.

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This event that occurs within the short normal polarity Jaramillo Subchron is correlated with warm Marine Isotope Stage 31 (Villa et al., 2007; Scherer et al., 2008). The anomalous warming implies an increase of 4–61C in SST, possible incursion of Subantarctic Surface Waters and depression of the lysocline – an event that is apparently unique in the last 3 Ma. The warming may occur during a unique conﬁguration in orbital parameters that may have produced an extended interval (20 kyr) of unusually warm (e.g. Raymo et al., 2006; Scherer et al., 2008), or alternatively unusually long (Denton and Huybers, 2008) SH summers. Planned future recovery of a 1–1.5 Ma ice core from Antarctica should shed valuable light on the greenhouse gas and temperature conditions at this time.

11.6.4. Mid-Pleistocene Climate Transition and Antarctica G-I variability in the 100 Kyr Glacial World The change in frequency of the G-I climate cycle between 900–700 kyr, from the 40 kyr duration, that dominated much of the glacial Cenozoic (last B35 Ma), to B100 kyr cycles remains one of the most poorly understood events in palaeoclimatology. Explanations have involved ampliﬁcation of the weak orbital inﬂuence by non-linear feedbacks in the carbon cycle (Shackleton, 2000), the internal dynamics of large continental ice sheets (Clark and Pollard, 1998), precessional-forcing (Raymo, 1997) and obliquity-beat skipping (Huybers and Wunsch, 2005). Depositional characteristics of the Late Pleistocene AND-1B G-I sedimentary cycles such as: the predominance of subglacial diamictites, thin interglacial ice shelf mudstones and no evidence for meltwater suggests a further cooling and stabilisation of the margins of the Antarctic Ice Sheets after the B-M polarity transition B800 kyr. A similar pattern occurs in ODP Site 1167 core from the Prydz Bay trough mouth fan which shows a signiﬁcant reduction in sedimentation rate and change in sediment provenance attributed to a less erosive, dry-based ice sheet from B1 Ma (O’Brien et al., 2007). The calving line of the Ross Ice Shelf appears not to have retreated south of its present interglacial position during subsequent ‘warmer-thanHolocene, super-interglacials’ (e.g. Marine Isotope Stages 11, 9 and 5; e.g. Jouzel et al., 2007). This evidence is in contradiction to hypothesised Late Quaternary collapse of the WAIS (Scherer et al., 1998) and far-ﬁeld evidence for sea-levels at þ20 m above present assigned to Marine Isotope Stage 11 (Hearty et al., 1999).

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Ocean records (e.g. ODP 1123 and 1090) imply northward expansion of seasonal sea ice, intensiﬁed thermohaline circulation and northward displacement of wind-driven, gyral circulation and subantarctic water masses from B700 kyr to present. The atmospheric temperature and greenhouse gas records from Antarctic ice cores show a pronounced 100 kyr periodicity that is coherent and in phase with marine temperature records (e.g. SSTs, Crundwell et al., 2008), but leads the deconvolved ice volume signal (e.g. d18O, Shackleton, 2000). The records of greenhouse gases from Dome C show very strong congruence with many features of the temperature record, and are consistent with CO2 in particular playing a signiﬁcant role in temperature ampliﬁcation, and also suggest that the Southern Ocean plays a leading role in controlling the atmospheric concentration of CO2 on G-I time scales. Most of the variation in total grounded ice volume in Antarctic glacial-cycle ice sheet models (e.g. Ritz et al., 2001; Huybrechts, 2002) is due to expansion and contraction of grounding lines across continental shelves, mostly in the WAIS Ross and Weddell sectors, but also to the west of the Antarctic Peninsula and other areas and is equivalent to B15–20 m of sea level. The EAIS interior responds in the opposite sense, contracting slightly at Northern Hemispheric glacial maxima due to lower model snowfall rates (which are reduced as temperatures fall). Glacial surface elevations based on ice core temperature modelling (Parrenin et al., 2007b) suggest central EAIS plateau decreased by B100–150 m compared to the present, with minor volume changes equivalent to just a few metres of sea level. In contrast ice build up on coastal Antarctica was of the order of several hundred metres during the LGM (e.g. Law Dome, Delmotte et al., 1999). In situ accumulation, rather than ﬂow, is indicated from the ice composition which precludes an interior Antarctic source. Higher dust ﬂux to Antarctica during glacial maxima compared with warm periods is considered to have radiative affects over Antarctica and provide nutrients (e.g. Fe) to the Southern Ocean promoting higher algal productivity and atmospheric CO2 drawdown (Wolff et al., 2006).

11.6.5. Millennial-Scale Climate Variability In general millennial-scale cycles occur in high-resolution Antarctic ice cores, particularly in methane and temperature records (e.g. Byrd, EPICA Dome C). However, they appear lower amplitude and their timing may differ from their NH counterparts (D-O events).

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The pattern of millennial-scale variability superimposed on G-I climate cycles in Antarctic ice cores occurs both during and prior to the last 100 kyr. The millennial-scale climate cycles labelled as ‘A-events’ in Antarctic temperature records suggest that Antarctica and the Southern Ocean may be out-of-phase with Greenland warming events (Blunier and Brook, 2001). It has been suggested that this was the pattern expected from simple models involving changes in ocean heat transport (Stocker and Johnsen, 2003), with Antarctica and the Southern Ocean able to accumulate heat when ocean heat transport to the north was weak. Although uncertainties between the age of the gas (methane) and the ice enclosing it in Antarctic ice cores makes evaluation of the precise timing of inter-hemispheric millennial-scale cycles difﬁcult, in some D-O cycles the temperature maxima in the north and south appear to be synchronous (e.g. EPICA Community Members, 2006). In this case, methane concentrations are most likely controlled by changing wetland extents in the northern high latitudes and tropics, along with changes in sinks (Valdes et al., 2005), and thus are relevant to Antarctic climate only as a relatively minor ampliﬁer. However, a Paciﬁc Basin ocean clathrate source for the methane has also been proposed (Kennett et al., 2000). Millennial-scale Antarctic warm periods (A1–A7; Blunier and Brook, 2001) are well expressed in SST, thermohaline circulation (sortable silt) and d18O records from the southern mid-latitude ocean (Barrows et al., 2007; Carter et al., 2008). One of these events, ACR, at B14.2–12.4 ka, is associated with an abrupt cooling of B21C in the Southern Ocean and an expansion of ice shelves and sea ice (Shemesh et al., 2002), modest intensiﬁcation of winds (Stenni et al., 2001; Ro¨thlisberger et al., 2002) and intensiﬁcation in deep abyssal inﬂow along eastern New Zealand through the Paciﬁc gateway (Carter et al., 2008). Thus, millennial-scale climate variability in Antarctica and Southern Ocean appears to display many similar characteristics and processes (perhaps with reduced amplitude) to those observed at the orbital scale (e.g. SSTs, sea ice, water masses and circulation).

ACKNOWLEDGEMENTS Referees comments by Phil O’Brien and Mike Bentley helped improve the text and sharpen some of the conclusions. Comments and edits on a draft manuscript by Robert McKay are greatly appreciated.

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Developments in Earth & Environmental Sciences, 8 F. Florindo and M. Siegert (Editors) r 2009 Elsevier B.V. All rights reserved DOI 10.1016/S1571-9197(08)00012-8

Chapter 12

Antarctica at the Last Glacial Maximum, Deglaciation and the Holocene A. P. Wright1,, D. A. White2, D. B. Gore2 and M. J. Siegert1 1

School of GeoSciences, Grant Institute, University of Edinburgh, The King’s Buildings, West Mains Road, Edinburgh, EH9 3JW 2 Department of Physical Geography, Macquarie University, Sydney, NSW 2109, Australia

ABSTRACT Recent technological advances in the study and dating of both land and marine glacial geologic features combined with improvements in both glaciological and post-glacial isostatic rebound modelling have led to signiﬁcant improvements in our knowledge and understanding of the Antarctic Ice Sheets at the Last Glacial Maximum (LGM) and their subsequent changes throughout the Holocene. Here we review the geological evidence for the extent and timing of the maximum advance of the East and West Antarctic Ice Sheets and the ice cover of the Antarctic Peninsula during the most recent glacial cycle. We also discuss evidence for the rate and timing of Holocene Ice Sheet retreat. Geological data provide a very important ‘ﬁrst-hand’ record of ice-sheet changes over a range of time periods. They are also useful for constraining and improving models which then have the potential to both ﬁll in the gaps for which geological data are unavailable, and to make predictions of the future. Inspection of the geological record allows us to form qualitative scenarios concerning glacial history. Numerical modelling has been used on several occasions to test such hypotheses. We discuss such numerical studies, indicating both their importance and limitations in order to develop quantitative ideas about the late Quaternary history of the ice sheet. An important environmental aspect of Antarctica’s glacial history is its contribution to global sea level rise since the LGM. The past decade has seen the range of estimates (from reconstructions based solely on geological evidence, on glaciological Corresponding author. Tel.: +44(0)131 650 7339;

E-mail: [email protected] (A.P. Wright).

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modelling constrained by geology and on modelling of the isostatic rebound) change from 0.5–38 m sea level equivalent to 5.9–19.2 m. Although the convergence of estimates is encouraging, there is a need for further glaciological modelling to take full account of the constraints placed on LGM ice expansion, particularly in East Antarctica, by new geological evidence.

12.1. Introduction Antarctica is comprised of two main grounded ice sheets; the West Antarctic and the East Antarctic Ice Sheets separated by the Transantarctic Mountains. Another signiﬁcant region of grounded ice exists in the Antarctic Peninsula where a series of ice caps and glaciers are located (Fig. 12.1). The present volume of ice in Antarctica is around 25.4 106 km3 (equivalent to 57 m of global sea level) of which 91% is within the East Antarctic Ice Sheet (Lythe et al., 2001). The maximum ice sheet surface elevation of 4093 m occurs at Dome Argus which, along with several other subsidiary ice domes connected by ridges, forms a major ice divide through the centre of East Antarctica. Ice thickness in the central regions of the continent varies between 2800 and 4500 m mainly as a consequence of bedrock topography which is known to vary spatially by more than a vertical kilometre over just a few km. If the ice were to be removed from East Antarctica, the bedrock surface would be largely above sea level. However, if the same were to happen over West Antarctica, even accounting for isostatic rebound, the bedrock would remain below the modern sea level. Thus, the West Antarctic Ice Sheet is often referred to as a marine-based ice sheet. Ice drains from the interior domes via fast ﬂowing rivers of ice known as ‘ice streams’, where the dominant method of ﬂow is by sliding or basal sediment deformation. These contribute grounded ice to numerous ﬂoating ‘ice shelves’ that surround the grounded ice continent. The Ross and the Filchner-Ronne ice shelves are Antarctica’s largest, with areas of 472,000 and 430,000 km2, respectively. Icebergs, usually of tabular form, calve from the marine margins of ice shelves, where ice thickness is usually 250–300 m. This process represents an important mechanism by which ice is lost from the ice-sheet system, accounting for 75–85% of all ice lost from Antarctica (Jacobs et al., 1992). A signiﬁcant amount is also lost due to bottom melting near the grounding lines of some outlet glaciers, particularly those where the grounding line is well below sea level (Rignot and Jacobs, 2002).

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Figure 12.1: Location map and surface elevation of Antarctica with place names mentioned in the text. Contours are at 500 m intervals.

12.2. Response of the Ice Sheets to Glacial Climate and Late Quaternary Ice-Sheet Reconstructions The climate around Antarctica during ‘‘full glacial’’ periods such as the LGM is likely to have been highly conducive to the presence of ice-sheets. Therefore, the ‘minimum’ reconstruction for the LGM Antarctic Ice Sheets is similar to the present geometry. However, due to the sea-level reduction of around 120 m (e.g., Bard et al., 1990; Shackleton, 2000) and ocean temperature changes that occurred during the LGM, the ice-sheets in Antarctica are known to have grown out towards the continental shelf edge in several places. The ice shelves bordering the Antarctic Ice Sheet thickened, grounded and became part of the parent ice mass. A ‘maximum’

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reconstruction would therefore be an Antarctic Ice Sheet which expanded to reach the continental shelf break right around the continent. There is some ice core evidence from around Antarctica, including Taylor Dome and Law Dome, which suggests a much lower rate of accumulation at the LGM. The Taylor Dome core also shows that LGM storm paths came from a different direction to the modern day. This suggests a reorganisation of the climate system at least at this part of the ice sheet (Morse et al., 1998). While reductions in both sea-level and air temperature at the LGM would have been conducive to ice-sheet expansion, decreasing accumulation may well have had a restraining effect on icesheet growth. To resolve how the ice sheet responded to such complicated changes in forcing requires the application of numerical models, which we will return to later. Ice-sheet behaviour in the late Quaternary is coupled with the cryosphere, ocean and atmosphere. For example, if sea ice extent increased due to cooling of the air temperature over the Southern Ocean (Armand, 2000), then the moisture supply to the ice sheet may have decreased and so ice-sheet growth may have been impeded. This introduces the possibility that independent mountain glaciers in Antarctica may have responded differently to LGM conditions than areas connected to an ice-sheet. In other words, ice-sheet advance may have been related to sea level fall while, at the same time, glacier decay may be related to a signiﬁcant reduction in rates of precipitation. Since the CLIMAP (1976) reconstruction of Ice Age Earth, there have been numerous reconstructions of LGM ice extent in Antarctica based on evidence from sediments and geomorphology, isostatic rebound and ice-ﬂow modelling. These reconstructions provide estimates ranging from 0.5–2 to 38 m of sea level equivalent (e.g., Budd and Smith, 1982; Nakada and Lambeck, 1988; Colhoun et al., 1992). While the techniques used vary, many of the early estimates placed the grounding line at the edge of the continental shelf, and have relatively steep ice proﬁles. As such these are at the upper end of estimates of sea-level lowering. More recent work has advanced our understanding of the dynamics of ice ﬂow (in particular the longitudinal proﬁle of ice streams and their controls), produced sophisticated 3D thermomechanical ice-sheet models and provided reliable geomorphic evidence that indicates relatively limited ice expansion in some areas. These advances in understanding have rendered the upper estimates implausible, such that recent reconstructions indicate an increase of Antarctic ice volume at the LGM with respect to today of only 5.9–19.2 m of sea level equivalent (Bentley, 1999; Huybrechts, 2002).

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12.3. Geological Information 12.3.1. Constraining Late Quaternary Ice-Sheet Extent, Volume and Timing The geometry of the Antarctic Ice Sheet throughout the last glacial cycle has been compiled using a number of different methods. In terrestrial environments, the aerial extent of former ice cover is generally recognised by mapping the lateral extent of ice marginal landforms such as moraines or ice marginal lake sediments, or by mapping subglacial landforms such as drumlins and striae. Ice volume can equally be reconstructed by mapping glacial debris striae, the locations of erratic boulders and the height/extent of mountain trimlines. The relative age of each mapped unit is usually constrained using a measure of exposure age, most often by the degree of weathering (such as soil formation or development of tafoni or iron staining on boulders), and more recently by measurement of isotopes produced by cosmogenic radiation. Stratigraphic correlations are generally difﬁcult due to lack of natural exposures. Similar techniques are also used in marine environments, but rather than investigating the geomorphology of former glacial landscapes using ﬁeld mapping and aerial photography, the ocean ﬂoor is imaged using techniques such as swath bathymetry and side-scan sonar. These techniques commonly identify ice marginal landforms such as moraines, or subglacial landforms such as mega scale glacial lineations, meltwater channels and drumlins. Stratigraphic techniques are more useful in the marine environment due to the ease of data acquisition via acoustic and seismic surveys, particularly when these sediment units are also sampled by gravity or piston cores. Establishing detailed late Quaternary glacial chronologies in Antarctica can be challenging (e.g., Anderson et al., 2002; Ingo´lfsson, 2004). The most direct dating is usually achieved in presently terrestrial environments, through cosmogenic exposure dating of moraines or radiocarbon dating of fossils contained within emergent marine sediments or ice-marginal lakes. However, factors such as production rate uncertainties (e.g., Gosse and Phillips, 2001), recycling of clasts with prior exposure and post-depositional reworking of glacial sediments (e.g., Brook et al., 1995) reduce the precision of cosmogenic chronologies to around 10% (Putkonen and Swanson, 2003). Where such materials are preserved, the algal mats commonly found in proglacial lake sediments can provide reasonably accurate radiocarbon ages, provided that the water column is well mixed and that at least part of the lake surface ice cover melts out each summer (Gore 1997b; Hendy and Hall, 2006). Limiting ages on glacial events in terrestrial environments can

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sometimes also be constrained by radiocarbon dating of the small amount of organic material sometimes found in ice-marginal or postglacial sediments (e.g., Burgess et al., 1994). Carbon sourced from marine environments, such as the bodies of seals and penguins or carbonate shells in former (and emergent) marine environments is subject to a marine reservoir effect that varies around the continent, but is on the order of 1.3 ka (e.g., Berkman and Forman, 1996; but also see Kiernan et al., 2003). On the continental shelf the majority of published data with regard to the timing of ice advance and retreat have been obtained by Atomic Mass Spectrometry (AMS) radiocarbon dating of the acid-insoluble (mostly diatom) fraction of organic material that is preserved immediately above or below a layer of glacially derived sediment. There are several potential problems with this technique, for example the inﬂux of ‘‘dead’’ carbon from ice melt, ﬁne-grained geogenic carbon such as graphite and recycling of older organic matter means that such material at the modern sediment/water interface provides radiocarbon ages of W2 ka, and sometimes W6 ka in areas of little biological productivity (e.g., Andrews et al., 1999). These values exceed the age derived from carbonate shells from living biota at such sites by over 1.3 ka. To some extent this problem has been rectiﬁed in more recent studies by AMS analysis of discrete carbonate shells (e.g., Leventer et al., 2006), but the lack of shells in many cores means that this approach is not always possible. Despite these difﬁculties, considerable efforts have been made in order to constrain the timing of ice advance and retreat across the continental shelf in many sectors of the continent. 12.3.2. Last Interglacial (120 ka BP) Data from a number of sources, including dated coral terraces and the oceanic d18O record, have ﬁxed the timing of the Eemian interglacial at 128– 116 ka BP (Stirling et al., 1998; Shackleton et al., 2003). During this period there was less ice in the world than at present and global sea level was between 3 and 6 m higher than today (Stirling et al., 1998; Overpeck et al., 2006). Some believe that this was caused by a smaller West Antarctic Ice Sheet (e.g., Mercer, 1978; Scherer et al., 1998), while others suggest the decay of the Greenland Ice Sheet was responsible (e.g., Cuffey and Marshall, 2000; Tarasov and Peltier, 2003). Evidence to support the reduction in volume of at least part of the West Antarctic Ice Sheet at the last interglacial comes from observations of marine diatoms of Quaternary age and the measurement of high concentrations of 10Be in sediments recovered from beneath the

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West Antarctic Ice Sheet (Scherer et al., 1998). 10Be and diatoms are thought to accumulate only under open marine conditions; when the sea is ice covered there is very little deposition of these materials. Moreover, if ice is grounded, any small amounts of 10Be and diatoms will be eroded. Sediments retrieved from beneath Ice Stream B, on the Siple Coast of the Ross Sea embayment, were probably emplaced during the last interglacial when marine conditions prevailed. The current dynamics of the ice-sheet may be affected by the location of these low shear strength, now subglacial, sediments (Anandakrishnan et al., 1998; Studinger et al., 2001). Similarly the Institute and Mo¨ller ice streams may have been free of grounded ice at the LGM, which according to Bingham and Siegert (2007) could account for B2 m of sea-level rise during the Eemian interglacial. 12.3.3. Last Glacial Maximum, the Holocene to the Present Day (B20–0 ka BP) Regardless of ice-sheet geometry at the last interglacial, there is persuasive evidence from the geological record to indicate that the Antarctic Ice Sheet was larger than present around the time of the global sea level lowstand at B20 ka BP, although the extent of this expansion is well constrained at only a few sites around the continental margin. Evidence of former ice surface elevations is particularly sparse, in part due to the lack of sites at which such information can be preserved, but also due to the difﬁculty in accessing remote inland mountain ranges where this evidence might occur. In the next sections we discuss glacial history by dividing Antarctica into seven sectors, based on lines of longitude (Fig. 12.2), which are approximately based on ice divides but centred on the geographical pole.

12.3.3.1. 0–601E: Queen Maud land/Enderby land Schirmacher Oasis (11.51E) is a 34 km2 area located behind Novolazarevskaya Ice Shelf. Infra-red stimulated luminescence dating of lake ﬂoor sediments yielded burial ages of B52 ka BP in Lake Glubokoye (Krause et al., 1997), which was in part corroborated by AMS 14C ages of 35 ka BP from 2 m above the glacial diamict at the base of the lake sediments. On the basis of these data, Schirmacher Oasis was ice free throughout the LGM. On the northern side of Fimbulheimen, the mountain range to the south of Shirmacher Oasis, deposition of mumijo (proventricular ejecta of the snow petrel, Pagodroma nivea) throughout the LGM indicates that ice thickening in

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Figure 12.2: Locations of the sectors discussed in this chapter; 1: 0–601E Queen Maud/Enderby Land. 2: 60–751E Mac.Robertson Land/Lambert Gl/Prydz Bay. 3: 75–1501E Princess Elizabeth/Queen Mary/Wilkes Land. 4: 1501E–1501W Transantarctic Mountains/Ross Sea Embayment. 5: 150–851W Marie Byrd Land/Ellsworth Land and the Amundsen and Bellingshausen Seas. 6: 85–601W Antarctic Peninsula/Bellingshausen Sea. 7: 60–01W Weddell Sea/Filchner Ice Shelf/Coats Land. this sector of the ice-sheet was also limited, with o80 m thickening occurring at the Insel Range (721S, 111E) during this time (Hiller et al., 1995). The well preserved 46–30 ka BP shorelines at Lutzow-Holm Bay (37.51E) indicate that this site probably remained ice-free through the LGM, constraining the expansion of the ice-sheet in this area (Igarashi et al., 1995; Igarashi et al., 1998). Holocene shorelines are also present at this site, reaching at least 15 m above present sea level and dating from 8 ka to the present day (Maemoku et al., 1997; Miura et al., 1998), suggesting a decrease in regional ice load in this area since the LGM (Nakada et al, 2000). Despite an abundance of small ice free areas that are suitable for Quaternary studies, including Riiser-Larsen (50.71E), Øygarden Group (57.51E) and Stillwell Hills (59.31E), there is a paucity of information regarding the geometry of this sector of the ice-sheet during or following the LGM.

12.3.3.2. 60–751E: Mac.Robertson Land/Lambert Glacier-Amery Ice Shelf/Prydz Bay 10

Be and 26Al cosmogenic isotope exposure ages on glacial erratics show that the ice-sheet thickness at Framnes Mountains (62.51E) was B350 m greater

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than present during the LGM, had begun lowering at 13 ka BP, and had reached the modern ice margin by 6 ka BP (Mackintosh et al., 2007). This evidence agrees reasonably well with side-scan sonar/piston core data, which indicate that ice had retreated from the mid-outer continental shelf in Nielsen Basin B80 km to the east of Framnes Mountains by B11 cal ka BP and had reached the inner shelf by 6 cal ka BP (Harris and O’Brien, 1998). Former ice advances of the Lambert Glacier-Amery Ice Shelf system (701E) into Prydz Bay (Fig. 12.3) are relatively well constrained through geophysical investigations and marine sediment cores (Domack et al., 1998; Taylor and McMinn, 2002), and by drilling through the ice shelf (Hemer and Harris, 2003). The sedimentary record indicates that ice was grounded on shallow (o500 m water depth) banks in Prydz Bay and across most of the area currently occupied by the Amery Ice Shelf at the LGM. However, the base of an uninterrupted sequence of glacial marine and biogenic sediments containing a relative abundance of organic material has been AMS radiocarbon dated to at least 30 ka BP within the Prydz Channel, a 700 m deep trench that cuts across the continental shelf in Prydz Bay (Domack et al., 1998). This suggests that grounded ice may not have extended all the way to the continental shelf at the LGM. Ages from acid-insoluble organic matter offshore from Amery Ice Shelf suggest a retreat timing of 11.5 cal ka BP for the Lambert Glacier-Amery Ice Shelf System (Domack et al., 1998). Investigations into glacial and lake sediments deposited on mountains ﬂanking the major outlet glaciers (Wagner et al., 2004; White and Hermichen, in press) also constrain the thickness of the ice sheet at the LGM. With the exception of the area around the southern tip of the modern Amery Ice Shelf, glacial sediments dating from the LGM are restricted to o200 m above the modern ice margin. Also, there are no emergent shorelines around epishelf Beaver Lake (Adamson et al., 1997), but there are subaerially deposited postglacial sediments on the lake bottom at 60 m below modern sea level (Wagner et al., 2007). These data support the results from the continental shelf, and indicate that both thickening and expansion of ice in this region at the LGM was limited. One reason for this may be the deep basin along which the outlet glaciers ﬂow, which reduces the ability of the ice-sheet to advance to the continental shelf edge (Taylor et al., 2004).

12.3.3.3. 75–1501E: Princess Elizabeth Land/Queen Mary Land/Wilkes Land Larsemann Hills (76.21E) hosts lake sediments dated to 14C background (W42 ka BP), and lake sediment stratigraphy which has been used to infer continual exposure since the last interglacial (Hodgson et al., 2005).

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An optically stimulated luminescence age of 21 ka BP was obtained from glacioﬂuvial sediments within 500 m of the ice margin (Hodgson et al., 2001), supporting interpretations of continual exposure through the LGM. Emergent shorelines are o3 m asl, consistent with relatively minor former ice loading. Vestfold Hills (781E) has emergent marine shorelines to o10 m asl (Zwartz et al., 1998), and cosmogenic 10Be analyses reveal that the ice margin had retreated to within 5 km of the present margin by 12.5–9 ka BP (Fabel et al., 1997). There has been a minor (o4 km) readvance of ice along the ﬂank of Sørsdal Glacier during the late Holocene (Adamson and Pickard, 1983; Gore, 1997a). Gaussberg (89.21E) is a 370 m high, glacially striated volcano on the coast. The benched morphology and presence of palagonite encrusted pillow lavas indicates that this eruption occurred in a water ﬁlled subglacial vault, and that the ice-sheet has since retreated to its present position since the eruption at 56 ka BP (Tingey et al., 1983). Bunger Hills (1011E) has emergent marine shorelines to o11 m asl (Colhoun and Adamson, 1992; Colhoun et al., 1992). Optically stimulated luminescence ages from glacial lake shorelines and glacioﬂuvial sediments indicate that deglaciation commenced B40–30 ka BP (Gore et al., 2001), with the area largely deglaciated by B25 ka BP. Like Vestfold Hills, the oasis attained its present form around 11 ka BP. These data are corroborated by 14 C ages indicating exposure through the LGM (Hedges et al., 1996, cited in Krause et al., 1997). Windmill Islands (110.31E) has emergent marine shorelines to 35 m asl (Goodwin, 1993; Goodwin and Zweck, 2000), with deglaciation of the southern islands by 11 cal ka BP and the northern peninsulas by 8 cal ka BP (Kirkup et al., 2002). Offshore from the large Mertz and Ninnis glaciers (145–1501E) of George V Land swath bathymetry has been used to identify mega-scale glacial Figure 12.3: (A) LGM ice-sheet surface reconstruction in Mac.Robertson Land and eastern Princess Elizabeth Land. Grey lines indicate present icesheet contours (in m asl), brown polygons indicate ice-free areas, and grey areas indicating ice-sheet ﬂow W300 ma 1 (Joughin, 2002). Solid green lines indicate LGM ice-sheet contours derived from ﬁeld studies, with investigated sites indicated by red circles (Domack et al., 1998; Harris and O’Brien, 1998; Zwartz et al., 1998; Hodgson et al., 2001; Taylor and McMinn, 2002; Whitehead et al., 2003; Leventer et al., 2006; K. Lilly, Pers. Comm., 2007; White and Hermichen, in press). Dashed lines indicate areas where heights or grounding lines are poorly known. (B) Proﬁle through the cross-section AB.

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lineations within the Mertz Trough which indicate LGM expansion of grounded ice to the outer continental shelf in this part of the ice-sheet (McMullen et al., 2006). Successive grounding zone wedge deposits have also been imaged in this region which record pauses in the retreat of the ice-sheet across the continental shelf, these however are yet to be conclusively dated (McMullen et al., 2006).

12.3.3.4. 1501E–1501W: Transantarctic Mountains/Ross Sea Embayment During the 1990s a comprehensive geophysical and sedimentological dataset was compiled across the Ross Sea using seismic proﬁles, swath bathymetry and side-scan sonar imagery of the sea-ﬂoor sediments in front of the present day Ross Ice Shelf (Anderson,1999; Domack et al., 1999; Shipp et al., 1999). Lineations, drumlins and large-scale grooves provide a ﬁrm basis for the identiﬁcation of former fast-ﬂowing, marine based ice streams situated in bathymetric troughs which extend right to the edge of the continental shelf in places but only to the mid-outer parts of the shelf in others (Fig. 12.4). Sediment cores taken from these mega-scale lineations have retrieved

Figure 12.4: LGM paleodrainage map for the Ross Sea based on geomorphic features on the shelf and including data from Shipp et al. (1999). Arrows are ﬂow directions based on lineations and the dashed line is the maximum grounding line position. Taken from Mosola and Anderson (2006), and reproduced with permission from Elsevier.

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unconsolidated Plio-Pleistocene strata that are interpreted to have composed the readily deformable bed which allowed such extended ice streams to develop (Mosola and Anderson, 2006). Lateral moraines deposited by the many glaciers ﬂowing through the Transantarctic Mountains and into the Ross Sea during the LGM have very similar proﬁles. These moraines are located high above the present day ice surface at the feet of these glaciers but merge toward the modern-day surface of the ice-sheet when traced upstream. LGM thickening indicated by these moraines therefore increases towards the marine margin, indicative of a thicker, grounded Ross Sea ice-sheet whereas the interior of the ice sheet in East Antarctica appears to have maintained a relatively constant thickness (Broecker and Denton, 1990). Deglaciation of the Ross Sea is thought to have begun ﬁrst at the western margin and proceeded like a ‘swinging gate’ hinged at the eastern side of the basin, with the LGM Ross Sea ice-sheet largely composed of extensions of the present day Siple Coast ice streams (Conway et al., 1999). The contribution of ice from East Antarctic glaciers, ﬂowing through the Transantarctic Mountains, to the Ross Sea ice-sheet remains contentious. The mineralogical composition of glacial diamicts from the ﬂoor of the eastern Ross Sea is similar to that of source areas on the Siple Coast, whereas diamicts from the western Ross Sea match with sources in the Transantarctic Mountains (Licht et al., 2005), indicating more complicated glacial ﬂow patterns at the LGM than a simple expansion of the present day Siple Coast ice streams (Fig. 12.5). Recent model investigations, which have attempted to reconstruct the late Holocene change in the thickness of Siple Dome from the depth-age relationship of the Siple Dome ice core and plausible scenarios for the changes in accumulation rate since the LGM, have led to estimates of surface height increase of 200–400 m. For the ice-sheet to have extended B1000 km to the continental shelf edge, the required ice surface gradient would have been much shallower than that of present day ice streams, and thus only possible by invoking a very slippery bed (Waddington et al., 2005). This supports the interpretation that ice from the Transantarctic Mountains must have contributed signiﬁcantly to the Ross Sea ice sheet. The timing of ice retreat in the eastern and central Ross Sea is poorly understood, but reworked foraminifera in glacial diamict on the outer shelf indicate that the maximum ice extent occurred after 16.5 cal ka BP. In contrast, there are numerous lines of both onshore and offshore evidence that constrain the timing of deglaciation in the western half of the Ross Sea. A recent review of the evidence from the marine record concluded that the maximum ice extent occurred after B17 cal ka BP (Licht, 2004),

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Figure 12.5: Flow lines for the Ross Ice Sheet during the last glaciation proposed by Licht et al. (2005) based on Ross Sea, East and West Antarctic till mineralogy and lithology. Dashed lines represent inferred ﬂow due to lack of sample coverage. Reproduced with permission from Elsevier. with deglaciation beginning by 14 cal ka BP (Domack et al., 1999; Shipp et al., 1999). The grounding line retreated relatively slowly along the western margin of the Ross Sea, reaching the modern ice shelf edge by 7 cal ka BP (Domack et al., 1999). This marine-based deglacial chronology is supported by evidence from terrestrial studies, including the timing of penguin recolonisation along the coast (Baroni and Orombelli, 1994), the altitude of large proglacial lakes in the McMurdo Dry Valleys dammed by an expanded Ross Sea ice sheet (Hall et al., 2002), cosmogenic exposure ages from moraines deposited by ice in the McMurdo Dry Valleys (Brook et al., 1995) and moraines and proglacial lake sediments on the Hatherton Glacier (Bockheim et al., 1989).

12.3.3.5. 150–851W: Marie Byrd Land/Ellsworth Land and the Amundsen and Bellingshausen Seas Exposure age dated moraines and recessional deposits indicate that ice thickness in Marie Byrd Land was around 45 m greater than present when deglaciation began around 10 ka BP (Ackert et al., 1999). Further surface

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exposure dating studies support the view that ice in this region was signiﬁcantly thicker at the LGM and that surface lowering has continued steadily throughout the Holocene (Sugden et al., 2006). Deglaciation in Marie Byrd Land therefore happened much later than the retreat of the Ross Sea grounding line and may still be in progress (Stone et al., 2003). Large scale troughs have been identiﬁed offshore from all the major present day outlet glaciers of this sector of West Antarctica, extending across the shelf of the Amundsen Sea (Anderson and Shipp, 2001) and the Bellingshausen Sea (O’Cofaigh et al., 2005b). Swath bathymetry has been used to determine that the grounded ice-sheet reached the shelf edge in outer Pine Island Bay (Lowe and Anderson, 2002; Evans et al., 2006, Fig. 12.6) where elongate subglacial bedforms in the soft sediments deposited at the base of a palaeo-ice stream end in gullies incised seaward of the shelf break (Dowdeswell et al., 2006). Likewise the Belgica Trough, which extends the full width of the continental shelf in the Bellingshausen Sea, exhibits geomorphological evidence of an ice stream which is thought to have drained much of Ellsworth Land as well as some of the southern part of the the Antarctic Peninsula Ice Sheet at the LGM (O’Cofaigh et al., 2005b). AMS dating of foraminifera near the Getz Ice Shelf (Marie Byrd Land) indicates that an ice stream in this trough retreated across a mid-shelf position by B13.4 cal ka BP (Anderson et al., 2002). The Pine Island Bay ice stream retreated from the mid-outer shelf after 17.574 cal ka BP, and had reached a position near the present day grounding line by 10.270.5 cal ka BP (Lowe and Anderson, 2002). Retreat from the shelf edge in the Bellingshausen Sea probably occurred at around the same time with a mid-shelf position reached by B14 cal ka BP at the latest (Pope and Anderson, 1992; Pudsey et al., 1994; Heroy and Anderson, 2005). The ice edge had retreated to the inner shelf at Palmer Deep by 13 cal ka BP (Domack et al., 2001), and there is no evidence for the presence of grounded ice on the mid-outer continental shelf after this time.

12.3.3.6. 85–601W: Antarctic Peninsula The late Pleistocene extent of grounded ice in the Antarctic Peninsula has received a signiﬁcant amount of attention during the past decade, partly due to ice-sheet models that predict a signiﬁcant increase in ice volume during this period (e.g., Huybrechts, 1990). Geomorphic evidence including glacial erratics and striations on nunatak summits along the length of the peninsula indicate that the ice sheet in this region thickened signiﬁcantly during the LGM. The centre of the ice-sheet thickened by up to 500 m in places,

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Figure 12.6: Geomorphological evidence for the extent and conﬁguration of the West Antarctic Ice Sheet in Pine Island Bay during the last glaciation (Evans et al., 2006). Also shown are the locations of postglacial iceberg scours (Lowe and Anderson, 2002; Evans et al., 2006). Reproduced by permission of Elsevier. reaching a maximum height of 2350 m asl at Mt Jackson (721S), and there is evidence for at least two distinct ice domes along the spine of the southern part of the peninsula (Bentley et al., 2000, 2006; Ingo´lfsson, 2004). The striation orientations in the south-western part of the peninsula also suggest

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that ice ﬂow was deﬂected by an expanded West Antarctic Ice Sheet in the Weddell Sea, indicating that ice advance in these two regions was synchronous (Bentley et al., 2006). Geophysical investigations of the continental shelf around the Antarctic Peninsula using a combination of multibeam swath bathymetry, side-scan sonar and seismic surveys have revealed a comprehensive set of subglacial landforms. The sites of former ice streams that drained the expanded Antarctic Peneinsula Ice Sheet at the LGM have been located at Marguerite Bay (O’Cofaigh et al., 2002, 2005a; Dowdeswell et al., 2004), in the western Bransﬁeld Basin to the north-west of the Peninsula (Canals et al., 2000, 2002; Evans et al., 2004) and the Robertson Trough on the eastern side of the Antarctic Peninsula (Evans et al., 2005) amongst others. In each case there is clear evidence that grounded ice streams extended all the way to the continental shelf break at the LGM. Post LGM retreat of the grounding line from the shelf edge position in the Robertson Trough region on the eastern side of the Antarctic Peninsula left no grounding zone wedges to indicate stillstands and is therefore interpreted to have been continuous and perhaps rapid (Evans et al., 2005). Slightly further north however, around the northern margin of the Larsen Ice Shelf A/southern part of the Prince Gustav Channel, ice retreat appears to have been much more gradual. The transition from subglacial to glacimarine deposition here has been shown by AMS radiocarbon dating to have occurred before 12 14C ka BP (Evans et al., 2005). A somewhat more complicated picture has emerged of the retreat history in the Marguerite Bay area but ﬁnal retreat across the continental shelf is known to have been underway by 13 14C ka BP (Pope and Anderson, 1992; O’Cofaigh et al., 2005a). Terrestrial and marine records are in broad agreement that the various presently ice-free coastal and bay regions of the Antarctic Peninsula were mostly deglaciated between 8 and 6 ka BP (Ingolfsson, 2004). The base of the glaciomarine sedimentary strata has been AMS radiocarbon dated to B8 ka BP in both the Gerlache Strait (Harden et al., 1992) and in Lallemand Fjord (Shevenell, et al., 1996). The terrestrial record includes fossil molluscs from raised marine deposits on King George Island that indicate deglaciation by around 9–8 ka BP (Ma¨usbacher, 1991) and glaciomarine and sub-littoral deposits overlying till on James Ross Island which began to be deposited around 7.4 ka BP (Hjort et al., 1997). The timing of ice-sheet lowering inland is less well constrained, but has been attempted using cosmogenic exposure dating. At George VI Sound, at Moutonne´e Valley (on Alexander Island) and in the Batterby Mountains on the western margin of the peninsula, ice downwasting may have begun as early as 25 ka BP and been complete by 15–10 ka BP (Bentley et al., 2006).

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12.3.3.7. 60–01W: Weddell Sea Embayment/Filchner-Ronne Ice Shelf/Coats Land Despite the potential of the Weddell Sea sector to have contributed a similar magnitude of post-LGM sea-level rise to that of the Ross Sea embayment, neither the extent of the grounding line or the change in thickness of the inland ice is yet well constrained (see Fig. 12.7). Glacial trimlines more than 1000 m above the present ice surface in the Ellsworth Mountains (Denton et al., 1992) are probably much older than the LGM (Sugden et al., 2006;

Figure 12.7: Ice-sheet reconstruction based on ﬁeld evidence from the Weddell Sea-Antarctic Peninsula region. Reproduced from Bentley et al. (2006) with permission from the GSA (see also Sugden et al., 2006).

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Fogwill et al., 2007). Cosmogenic exposure-age dating of surface bedrock in the Shackleton Range has also been used to constrain the LGM thickening of the ice near the present day at grounding line of the Filchner Ice Shelf to a maximum of 750 m and a more probable limit of 340 m, consistent with limited LGM thickening (Fogwill et al., 2004). Marine geological evidence for the location of the LGM grounding line in the Weddell Sea is perhaps even more sparse than terrestrial data of ice-sheet thickness. Glacial diamict sampled from the bed of Crary Trough, which underlies the Filchner-Ronne Ice Shelf and extends out towards the edge of the continental shelf east of Berkner Island, has a composition indicative of a source in West Antarctica (Anderson et al., 1991). Radiocarbon dating, however, has so far been unable to conﬁrm the age of these deposits. 12.3.4. Holocene Glacier and Climate Fluctuations Evidence for smaller ice volumes/ice extents during the mid-late Holocene relative to the present day is available at a number of sites. Sediments from the bed of epishelf Moutone´e Lake indicate the free movement of icebergs, and therefore the collapse of the George VI Ice Shelf between Alexander Island and the Antarctic Peninsula, between 9.6 and 8 cal. ka B.P (Bentley et al., 2005). Likewise the analysis of sediment cores recovered from the seabed of the Prince Gustav Channel on the eastern side of the peninsula, has determined that between 5 and 2 ka BP clasts were sourced from a range of distal locations, indicating open marine conditions and the collapse of the Prince Gustav Channel Ice Shelf (Pudsey and Evans, 2001). Neither is the mid-Holocene minimum conﬁned to the Antarctic Peninsula. Goodwin (1996) provides evidence from a number of sources that Law Dome was at least 3–4 km smaller before B4 ka BP, while Baroni and Orombelli (1994) show that ice shelves in Terra Nova Bay were reduced by a similar amount between 7.5 and 5 ka BP. On the contrary, there is geological evidence to suggest that the Larsen Ice Shelf B had been in place for W11 ka before its collapse in 2002 (Domack et al., 2005). In East Antarctica, Verleyen et al. (2005) provide relative sea level (RSL) based evidence for increased ice load near the Larsemann Hills between B7 and 2.5 ka BP. Given the similarity of the RSL curve to that obtained from the Vestfold Hills (Zwartz et al., 1998) this may have been more than a local feature. This ﬂuctuation may correlate with the glacial advances identiﬁed through striae patterns and weathering in the Vestfold Hills (Adamson and Pickard, 1986) and in the Rauer Group (White et al., submitted).

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12.3.5. Discussion: Pattern and Timing of Ice Retreat The geological record provides convincing evidence that the Antarctic Ice Sheet was more extensive at the LGM and has subsequently reduced in area and thinned near the margins. The magnitude of ice retreat varies around the continent, ranging from relatively little marginal change at some terrestrial sites such as the Larsemann Hills, to retreat of the grounding line by hundreds of km in the large embayments currently occupied by the Ross, Filchner-Ronne and Amery ice shelves. At most of the marine sites studied, the evidence implies that the LGM maximum ice extent was relatively shortlived, while some of the terrestrial sites (e.g., Bunger Hills, Lutzow-Holm Bay) suggest that the LGM was not necessarily the period of greatest ice extent during the last glacial cycle, and that ice-sheet margins were more advanced prior to 35 ka BP. Similarly, evidence for the timing of ice retreat following the LGM supports different retreat timings at different sites (Anderson, 1999). Some areas, such as the Antarctic Peninsula and the Bellingshausen Sea appear to have responded relatively rapidly to global climate and sea level changes following the LGM. At other locations, including the Ross Sea and the East Antarctic margin at Framnes Mountains, ice retreat appears to have begun relatively late and, in some places such as Marie Byrd Land, appears to have continued to the present day. While the evidence for this differential retreat timing cannot yet be considered deﬁnitive, the pattern may point to the differing sensitivities of each region to external forcing.

12.4. Numerical Modelling Reconstructions 12.4.1. Ice-Sheet Modelling Recent interest in the response of the Antarctic Ice Sheet to changing climates has been an important factor in driving the development of improved glaciological ice-sheet models. As well as three-dimensional ice ﬂow by internal deformation, these models now incorporate a temperature calculation throughout the ice sheet. Thermo-mechanical coupling allows a more realistic representation of ice ﬂow with the additional beneﬁt that basal sliding can be parameterised over areas of the bed that reach the pressure melting point. Ice-sheet models require inputs of bedrock elevation, initial ice thickness, sea level and climate. They can be used to provide quantitative information about ice ﬂow and the response of an ice sheet to changes in the

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model parameter values or to a change in external forcing. Models are best used along with ﬁeld data – both present day ice-sheet extent and past extents inferred from geological evidence. The ability of models to make predictions in areas of sparse data and for time periods which cannot be constrained by geological evidence means they are considered a valuable tool in reconstructing past ice-sheet conﬁgurations. The ﬁrst comprehensive glaciological modelling of the Antarctic Ice Sheet over a full glacial-interglacial cycle was published by Huybrechts (1990). This work has now been updated with several improvements to both the model and the climatic input (Huybrechts, 2002). An independent study with an alternative model formulation by Ritz et al. (2001) provides a useful comparison for the case of reconstructing the extent and thickness of Antarctica at the LGM. This model has also recently been forced with output from a Global Climate Model providing a further comparison (Philippon et al., 2006). Both Huybrechts (2002) and Ritz et al. (2001) use time-dependent climatic input to force their models. In both studies a reference simulation is produced using paleoclimatic information derived from the Vostok ice core (Petit et al., 1999) to run the model through four glacial-interglacial cycles (B400 ka). Free parameters in each model are then tuned to match the results to observations of the present day ice-sheet. The ice-sheet produced at the end of the reference run is then used as the initial condition for a range of further simulations to assess the model sensitivity to changes in parameter values and forcings. 12.4.2. Reconstructions of the Present Day/Interglacial Ice Sheet Both model representations of the present day are similar, with both showing some differences from the observed geometry, mainly around the margins. Both models over-predict the extent of grounded ice in the Amundsen Sea and the Amery Ice Shelf regions. Ritz et al. (2001) also over-estimate the extent of grounded ice in the Weddell Sea, while both models underestimate the extent of grounded ice in the Ross Sea (Fig. 12.8A, B). Ritz et al. (2001) ascribe the discrepancies between their model and present day observations to the coarse (40 km) grid resolution of their model, as this hinders their ability to reproduce fast ﬂowing ice streams and ﬂow retarding bedrock features. The result is an ice-sheet which covers a similar area but has a 12% larger volume. Overall the present day ice sheet produced by Huybrechts (2002) is a better match to the observations and this is probably due to the higher resolution (20 km) grid cells of that model.

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Figure 12.8: Numerical model reconstructions of the present day ice sheet (A, B) and of the Last Glacial Maximum (C, D) by Huybrechts (2002) and Ritz et al. (2001), respectively. Reproduced with permission from Elsevier and the AGU. 12.4.3. Time Dependent Simulations Both Huybrechts and Ritz run their models over 400,000 years encompassing the last four glacial-interglacial cycles. The forcing in each case consists of time-series of air temperature change and eustatic sea-level variation. Both studies use temperature records derived from the dD (deuterium) signal of the Vostok ice core. In each case accumulation is based on the present day

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pattern with the magnitude of variation calculated as a function of air temperature by the method described in Lorius et al. (1985). Huybrechts uses the SPECMAP reconstruction of sea-level change while Ritz et al. prefer that of Bassinot et al. (1994); the two records have some differences but sensitivity studies by Ritz et al. suggest that this does not affect their model results. Overall the two reconstructions agree on several important aspects of Quaternary Antarctic evolution: The present day ice sheet is slightly larger than the minimum reached during the last interglacial. Ritz suggests that this is because the present ice sheet is still in decline, while Huybrechts determines that the ice-sheet retreat from the Weddell Sea was more advanced during the last interglacial. General responses that dominate during a glacial period are grounding line advance and interior ice growth in West Antarctica with surface lowering and little margin advance in East Antarctica. Eustatic sea-level variation is the forcing factor driving change in the position of the ice sheet grounding line while changes in temperature (and consequently rate of accumulation) have the largest effect on surface altitude in the interior of East Antarctica. Grounding line advance occurs more readily in the Weddell Sea embayment than in the Ross Sea. Both models cite the differences in bed topography – the depth of the Ross Sea increases more rapidly with distance from the present coastline – as the cause of this asymmetry. The patterns of ice ﬂow and positions of the ice domes in East Antarctica are not signiﬁcantly changed from their present conﬁguration at the LGM. Ice ﬂow in West Antarctica would have been very different due to the advanced margins and thicker inland ice. There are several points at which results from the two models diverge. Huybrechts (2002) puts the timing of the LGM in Antarctica at 15 ka BP when the ice sheets would have contained 19.2 m sea level equivalent greater mass than at present, equivalent to an increase of 5 m sea level equivalent on the sea level contribution predicted with the same model in an earlier study (Huybrechts, 1991). Ritz et al. (2001) place the LGM at around 18 ka BP and ﬁnd a difference in mass of just 5.9 m sea level equivalent from today. The Ritz model generally has a lower surface slope around the margins due to its inclusion of a grounded ‘dragging shelf ’ transition zone between inland and ice shelf ﬂow intended to represent ice streaming. This factor is probably responsible for the majority of the differences between the two models, with the Ritz model placing thinner ice over much of West Antarctica at the

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LGM, especially in the Ross Sea and also notably in the present day area of Prydz Bay in East Antarctica (c.f. Fig. 12.8C and D). Interestingly, the reconstruction of Ritz et al. (2001) fails to predict any signiﬁcant expansion of ice covering the Antarctic Peninsula except in the very southern part where it joins the West Antarctic Ice Sheet (Fig. 12.8D). This is in contrast to the results consistently obtained with Huybrechts model, which in agreement with geological evidence outlined earlier predicts an expansion of the Antarctic Peninsula Ice Sheet to the edge of the continental shelf under LGM conditions (Fig. 12.8C). 12.4.4. Ice-Sheet Sensitivity Both Huybrechts and Ritz analysed the sensitivity of their respective models to variations in climatic forcing and to uncertainty of the parameter values by repeating the time dependent simulations described above and comparing snapshots of the ice-sheet at ﬁxed time-slices. These experiments dealt with aspects of the ice-sheet dynamics including bedrock adjustment, thermodynamic coupling, basal sliding and the response of ice shelves to ocean warming as well as to variations in air temperature and accumulation forcing. The elements of sensitivity common to both models include: The ice-sheet volume is relatively sensitive to changes in LGM snow accumulation. Ritz et al. found that reducing glacial accumulation values to 40% of present day (c.f. 50% for the standard simulation) decreased the area of the LGM ice-sheet by 18% and the volume by 43%. Grounding line advance during glaciations is sensitive to the way ice shelves respond to temperature change. Not allowing the ice temperature (which controls viscosity) to cool and hence stiffen the ice or preventing the basal melt rate from decreasing signiﬁcantly reduces the advancement of the grounding line during the LGM especially in West Antarctica. Sea level lowering appears to be necessary for grounding line advance but so does a decrease in surface temperature. The rate of crustal response appears to have an intricate relationship with the amount of grounding line retreat occurring at the glacial-interglacial transition. A more rapid isostatic response leads to earlier re-grounding of a retreating ice shelf in either the Ross or Weddell Sea embayments thereby reducing the total retreat. Finally, a lower value for the geothermal heat ﬂux reduces the temperature of basal ice and limits the area of the base reaching the pressure melting point. The colder ice ﬂows more slowly and the result is to increase ice-sheet thicknesses at the LGM.

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12.4.5. Recent Modelling of Antarctica Recent studies using a variety of dating techniques have indicated that the former extent of the East Antarctic Ice Sheet at the LGM may have been less than previously thought (e.g., Gore et al., 2001; Hodgson et al., 2001; Mackintosh et al., 2007). Past modelling studies however, including both those by Ritz and Huybrechts described above, have placed the margin of the East Antarctic Ice Sheet at the edge of the continental shelf in agreement with the CLIMAP reconstruction (CLIMAP, 1976). With this in mind some new numerical modelling experiments have been undertaken in order to investigate LGM ice sheet conﬁgurations compatible with the newly available geological evidence from East Antarctica. The GLIMMER community ice-sheet model is an implementation of the shallow-ice approximation with full thermodynamic coupling and the inclusion of basal sliding (Payne, 1999). Using this model an ice sheet was created on a slightly modiﬁed and isostatically rebounded BEDMAP topography (Lythe et al., 2001) with 20 km grid resolution. Model parameter values were tuned to reproduce as closely as possible the present day extent of the Lambert Graben region of the East Antarctic Ice Sheet (Fig. 12.9A). The accumulation rate was calculated for each grid cell as a function of continentality (proximity to the coastline) following the method of Oerlemans (1982); surface air temperature was calculated using a ﬁxed lapse rate and mass loss at the marine margins as a prescribed calving percentage. Table 12.1 shows the parameter values supplied to the model for the simulations of both the present day and LGM ice sheets. For the LGM these values are taken from the general consensus of the paleoclimate literature (e.g., Petit et al., 1999; Siegert, 2003; Van Ommen et al., 2004). The extent and thickness of the present day ice sheet produced by the model is a good match with observations in the Lambert Glacier region of East Antarctica (c.f. Fig. 12.9A with Fig. 12.3). The shape of the ice-sheet around the Lambert Graben and the surface height of the interior ice-sheet are reasonably well reproduced; the model used here is not capable of reproducing ice shelves and hence the Amery Ice Shelf does not appear in the results. A characteristic of this reconstruction of the LGM ice sheet (Fig. 12.9B) is the uneven nature of the glacial advance around the margins. The location of the grounding line at the LGM is predicted by the model to have advanced B80 km towards the middle of the continental shelf on the western ﬂank of the Lambert Graben, by only B40 km on the eastern ﬂank and for no signiﬁcant advance on the present day margin along much of the rest of the coastline (Fig. 12.9B, C). This pattern is matched by negligible changes in

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Figure 12.9: Numerical model results for the Princess Elizabeth – Lambert Graben – Mac.Robertson Land area of the East Antarctic Ice Sheet. (A) Reference experiment created with parameter values tuned for the present day. (B) Ice sheet reconstruction with best-estimate parameter values for the LGM. (C) Predicted ice surface height change (B-A) at the LGM. The locations of the two longitudinal proﬁles in Fig. 12.10 are marked A–B and C–D.

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Table 12.1: Values of model parameters used for the simulation of both the present day and LGM East Antarctic Ice Sheet.

Sea level air temperature (1C) Air temperature lapse rate (1C m 1) Relative precipitation Eustatic sea level (m) Calving fraction at marine margins Basal traction factor Geothermal heat ﬂux (W m 2)

Present day (0 ka BP)

LGM (B21 ka BP)

15 0.008 100% 0.0 90% 1.0 10 4 56 10 3

25 0.008 50% 120.0 90% 1.0 10 4 56 10 3

ice-sheet thickness along parts of the coast which are coincident with signiﬁcant thickening in other areas (Fig. 12.9C). The patterns of advance and thickening seen here are strongly inﬂuenced by the bathymetry of the continental shelf. Where the shelf is shallow the 120 m sea-level drop at the LGM allows expansion of the ice sheet. Under this scenario the extent of the ice-sheet is controlled largely by eustatic sea-level and hence by the volume of northern hemisphere ice sheets (e.g., Denton et al., 1986). In contrast the interior of the ice-sheet experiences a general decrease in thickness of between 100 and 150 m under the LGM conditions prescribed here (Fig. 12.9C). This is a result of the decrease in accumulation which is not mitigated by the decrease in air temperature as surface melting is relatively unimportant here under either present day or interglacial conditions. The locations of the two proﬁles shown in Fig. 12.10 were chosen to allow direct comparisons to be made with the LGM reconstructions of the icesheet surface from cosmogenic dating of glacially abraded surfaces and transported clasts (Fig. 12.3 and Mackintosh et al., 2007). The AB-proﬁle shown in Fig. 12.10A follows the route of the Fisher Glacier to its conﬂuence with the Lambert Glacier and where the continuation of the latter discharges into the Amery Ice Shelf. Though the model is not capable of reproducing ice shelves, it can be seen (by comparing Figs. 12.3B and 12.10A) that grounded ice in the model extends B100 km past the position of the grounding line inferred from GPS measurements of tidal motion (Fricker et al., 2002). This over-advance under present day conditions occurs possibly because the depth of the topographic low, near the deepest point of which is located the present grounding zone, is underestimated in BEDMAP. The present day model terminus position of the glacier in the Lambert Graben is in fact very similar to the location of the LGM transition zone between steady icesheet ﬂow and ice streaming inferred from geological evidence (Fig. 12.3B).

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Figure 12.10: Longitudinal proﬁles through the model reconstructed ice sheet for Fisher-Lambert Glacier and Framnes Mountains. These proﬁles demonstrate the main features of the LGM reconstruction of East Antarctica i.e. surface lowering in the interior with little or no advance of the margins. It is therefore perhaps unsurprising that no further terminus advance occurs in this area when the model is run under LGM climate/sea-level conditions (Fig. 12.10A). The LGM advance of the East Antarctic Ice Sheet margin along the CD proﬁle (Fig. 12.9) has been shown by Mackintosh et al. (2007) to have been only around 5–10 km.This is beneath the grid resolution of the present model and hence the unchanged margin position along the CD-proﬁle under LGM conditions (Fig. 12.10B) is in agreement with the ﬁeld evidence. The model also predicts a slight LGM lowering of the ice-sheet surface along the entire length of this proﬁle as a result of decreased accumulation, counter to the geological evidence which suggests a moderate increase in thickness near the margin (Mackintosh et al., 2007). The spatial extent of this thickening, however, is uncertain and may be too small to be resolved by the current model.

12.4.6. Isostatic Modelling of Post-Glacial Uplift and Contribution to Global Sea Level The Earth’s crust is displaced by the loading of ice sheets, which recovers when these loads are taken away. In Antarctica, the pattern of recovery

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following removal of an ice load has been reconstructed by investigating and dating records of relative sea level, such as of emergent shorelines (e.g., Adamson and Pickard, 1986; Baroni and Orombelli, 1994), the transition from marine to freshwater sediment deposition in enclosed basins (Goodwin, 1993; Zwartz et al., 1998; Verleyen et al, 2005) and by determining the present day rate of crustal motion by measuring the change in height of bedrock surfaces over time using a Global Positioning System (GPS; e.g., Tregoning et al., 2000). These uplift data have been compared with the response predicted by numerical Earth models to constrain the spatial and temporal pattern of former ice loads. However, some of the parameters (such as mantle viscosity) used in the models are not independently known, and it is often possible to model the measured isostatic response at any one site with a range of former ice loads (e.g., Bassett et al., 2007). Thus, care must be taken when interpreting these results in regions like Antarctica where relative sea level reconstructions are few and far between. Despite these issues, the technique is useful in constraining former ice heights where evidence is ambiguous, or difﬁcult to reconstruct by other means. This has resulted in numerous reconstructions of former ice heights across the continent (Nakada and Lambeck 1988; Tushingham and Peltier, 1991; Colhoun et al., 1992; Peltier, 1994), which provide a very large range of post-LGM sea level contributions from Antarctica, of between 2 and 37 m. The more recent models, which incorporate a greater degree of constraints derived from the geological studies detailed above, provide estimates of 17.3 m (ICE-5G model of Peltier, 2004) and 6.6–16.7 m (ANT5 and ANT6 of Nakada et al., 2000). These models have also been used to provide estimates of former ice loads at individual sites around the continent. Zwartz et al. (1998) used relative sea-level data from the Vestfold Hills to estimate that the ice sheet retreated 30–40 km, and thinned 600–700 m in this region since the LGM. Goodwin and Zweck (2000) used uplift information from raised beaches and the marine-lacustrine sediment transition in basins in the Windmill Islands and Budd Coast as input to an isostatic model of the Law Dome ice cap in East Antarctica. Their numerical Earth model predicts that the area near the present day margin of Law Dome formerly hosted an ice load of between 770 and 1000 m, and that ice expanded out 40 to 65 km across the continental shelf (Goodwin and Zweck, 2000). Goodwin and Zweck (2000) also used the similarity in the height of the marine limit along the coast between Wilkes and Oates land to postulate that the East Antarctic Ice Sheet in this area has retreated by a similar amount since the LGM.

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12.5. Summary There is persuasive marine geological evidence from the relatively well studied Ross Sea embayment, the Bellingshausen Sea, the Amundsen Sea and off the west coast of the Antarctic Peninsula for an LGM ice-sheet grounding line at, or very near to the edge of the continental shelf. A similar situation most likely existed on the eastern side of the peninsula and in the Weddell Sea embayment, though this region has been less well studied to date. LGM thickening of the West Antarctic Ice Sheet has been identiﬁed in several locations with the ice surface in the central regions likely to have been several hundred metres higher than today. Likewise ice in the Antarctic Peninsula can be shown to have been up to 500 m thicker at central sites, with the present day ice caps and glaciers coalescing to form the Antarctic Peninsula Ice Sheet at that time. The smaller number of studies undertaken on the East Antarctic Ice Sheet have found wide variations in the extent of LGM ice advance across the continental shelf. It appears that the grounding line in East Antarctica advanced to the continental shelf break in places, to a mid shelf position in others and for no signiﬁcant advance to have occurred elsewhere. Furthermore there are several sites, including the Bunger Hills, where deglaciation appears to have begun at around 30 ka BP, well before the LGM. The timing and rate of deglaciation during the Holocene appears to have varied greatly across the continent. The Antarctic Peninsula Ice Sheet and the Bellingshausen Sea sector of the West Antarctic Ice Sheet may have been amongst the ﬁrst regions to reach their present conﬁgurations while the Ross Sea sector along with many coastal areas in East Antarctica have responded less rapidly to climate and sea-level change since the LGM. At the end of the 20th century estimates of the reduction in the total volume of the Antarctic Ice Sheets since their maximum during the last glacial cycle ranged from 0.5 to 37 metres of equivalent sea level rise. Recent analysis of the geological evidence has made the upper estimates seem improbable with the consensus of opinion tending towards the upper end of the lower half of this range. Though there is as yet no single agreed value, recent estimates based on geological evidence (Bentley, 1999, 6.1–13.1 m; Denton and Hughes, 2002, 14 m) have converged with those based on numerical ice-sheet modelling (Ritz et al., 2001, 5.9 m; Huybrechts, 2002, 19.2 m; Philippon et al., 2006, 9.5–17.5 m) and those based on modelling of isostatic post-glacial uplift (Nakada et al., 2000, 6.6–16.7 m; Peltier, 2004, 17.3 m) to narrow the likely range to 5.9–19.2 m sea level equivalent.

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Subject Index AABW (see Antarctic bottom water (AABW)) AAIW (see Antarctic intermediate water (AAIW)) ACC (see Antarctic circumpolar current (ACC)) ACE programme (see Antarctic Climate Evolution (ACE) programme) Acoustic basement rocks, of Ross Sea region, 122 Acoustic stratiﬁcation, wedge downlaps, 145 Acoustic stratigraphy, of Wilkes Land, 127–133 Age-depth model, for ODP drilling site, 141 Alkalic magmatic province, Ross Sea tectonic evolution, 283 AMS (see Atomic mass spectrometry (AMS)) Andean Orogen, Antarctic, 275 ANDRILL drilling, on Ross Ice shelf, 124 Angiosperms, fossil ﬂora, 333–335 Antarctica climate change/warming, 103–107 future geoscience investigations in, 287–291 modelling, 555–558 Antarctic Andean Orogen, 275 Antarctic bottom water (AABW), 94–97

Antarctic circumpolar current (ACC), 97–99, 341, 573 Antarctic climate glacial-interglacial time scales, and, 481–485 millennial-scale changes, and, 485–486 Antarctic climate evolution, 509–513 ice sheet expansion, 509 late Pliocene cooling, 509 late Pliocene-early Pleistocene G-I climate cycles, 509–510 marine ice sheet margin development, 509 Marine Isotope Stage 31, 510–511 mid-Pleistocene climate transition, 511–512 millennial-scale climate variability, 512–513 Antarctic Climate Evolution (ACE) programme, 2–4 Antarctic coastal and continental shelf rock-drilling sites, 60–61 Antarctic coastal current, 100 Antarctic geological drilling programme (ANDRILL), 476–479 Antarctic glacial history, Cenozoic, 4–9 advancement in geological research during twenty-ﬁrst century, 58–63 developments in late 1970s, 46–47

578

Subject Index

discoveries at offshore and on continent in 1980s, 47–53 drilling records during 1972–1975, 39–45 land-based drilling, 1973–1974, 43 sea ice drilling, 1975, 43–45 ship-based drilling, 1972–1973, 40–43 events in evolution of, 36–37 research advancement in mid-twentieth century, 38–39 research advancement in 1990s, 53–57 Antarctic glaciological data versus Ice-Sheet models, 245–249 Antarctic Ice Sheet DSDP–ODP drilling dates initiation of, 50–52 dynamic, late Oligocene, 49–50 future prospective of, 63–69 John Mercer’s hypothesis for behaviour of, 47 landscape modelling, 58–59 modelling of, 55–56 numerical modelling of, 235–253 EISMINT, 242–245 ice-sheet models, 238–241 ice-sheet models versus glaciological data, 245–249 ice-sheet processes, 236–238 ice-sheet reconstructions, 249–253 model input, 241–242 Pliocene record of, 7 pre-Quaternary history for, 39 proxies climate records, 50 thermal conditions, 247–248 velocity map of, 68 Antarctic Ice Sheet, 468, 533–535 over a full glacial-interglacial cycle, 551

Antarctic intermediate water (AAIW), 90–92, 491 Antarctic Isotopic Maxima (AIM), 486 Antarctic Offshore Stratigraphy (ANTOSTRAT) project, 52–53, 116 and ODP, 53–54 Antarctic Peninsula, 472–473, 532, 545–547, F12.1 eastern margin of bathymetric contours, 153 single-channel seismic line across, 154 fossil ﬂoras King George Island, 328–330 Late Palaeocene-Late Eocene, 326–327 Seymour Island, 330 geologic transition in, 170 geomorphic and seismic stratigraphic features of, 162–164 lithostratigraphic transitions in, 167–170 Paciﬁc margin of depositional sequences, 159 mid-shelf basins, 161 ridge-crest segments, 158 sea-level and ice-volume changes, 170–172 sedimentary strata, 311–313 seismic stratigraphic history Cretaceous to early Oligocene, 165 Oligocene glaciomarine sections, 166 sediment drape, 167 South Orkney Islands region, 155–156 South Shetland Islands region Bransﬁeld Strait, 156–157 scientiﬁc drilling in, 158

Subject Index

terrestrial environment, Middle Miocene to Pliocene, 417–421 Antarctic periphery climate, 486–487 Antarctic sea-ice drilling, 54–55 Antarctic slope front, 100 ANTOSTRAT (see Antarctic Offshore Stratigraphy (ANTOSTRAT) project) Apparent sea level, 379 Atlantic DWBC inﬂow, 101–102 Atmospheric carbon dioxide (CO2) at E/O boundary, computer modelling, 344–348 inﬂuence across Oligocene–Miocene boundary, 385 Atmospheric circulation, 467 Atmospheric variability, from ice cores, 479–488 Antarctic climate and glacial-interglacial time scales, 481–485 Antarctic periphery climate, 486–487 future challenges, for ice coring, 487–488 ice core record, 480–481 millennial-scale changes, 485–486 Atomic mass spectrometry (AMS), 536 Austro-Hungarian North Pole Expedition, 15–16 Basement rocks, acoustic of Ross Sea region, 122 Battye Glacier Formation, 412 Beacon Supergroup, 273–274 Beaver Lake, 539 BEDMAP topography, 555 Bellingshausen Sea Volcanic Group, 421 Biomagnetostratigraphic dating, 55

579

Borchgrevink orogeny, 272–273 Bransﬁeld Strait, 156–157, 276 Bunger Hills, 541 Byrd, Rear-Admiral Richard E., 21–22 Byrd glacier, 248 Canonical Milankovitch theory, 467 Cape Roberts drilling project, 291 Cape Roberts Project (CRP), 54–55, 469 Carbon dioxide (CO2) (see Atmospheric carbon dioxide (CO2)) CDW (see Circumpolar deep water (CDW)) Cenozoic geological time periods, 2 magmatism in Ross Sea, during, 279–284 oxygen isotope records for, 343 tectonic evolution in Ross Sea, during, 279–284 Cenozoic climate history Antarctic Peninsula, 152–161 Prydz Bay, 135–144 Ross Sea, 118–126 Weddell Sea, 144–152 Wilkes Land, 126–135 Cenozoic investigations in the western Ross Sea (CIROS) Project, 47–48, 469 Circum-Antarctic geographical distribution maps dinocysts, in Late Paleocene–Miocene, 338 Circum East Antarctic Mobile Belt, 263 Circumpolar deep water (CDW), 92–94 types of, 94

580

Subject Index

CIROS (see Cenozoic investigations in the western Ross Sea (CIROS) Project) CIROS-1, 49–50 CIROS-1 drill core climate records, Oligocene–Miocene boundary, 381, 384–385 revised age models for, 383 western Ross Sea, 317–321 CLIMAP reconstruction, 555 Climate change/warming, of Antarctica and Southern Ocean (Subpolar region), 103–107 Climate modelling, 55–56 Climate records drilling, 316–321 Eocene/Oligocene, 310–349 from fossil plants, 325–331 from marine microfossils, 335–342 from palynomorphs, 331–335 from sedimentary record, 311–325 from terrestrial realm, 325–335 Middle Miocene to Pliocene marine record, 424–436 terrestrial environments, 404–424 Oligocene–Miocene boundary, 370–391 Coastal current, Antarctic, 100 Coats Land, 548–549 CO2 concentrations, 483 Collins Glacier deposit, fossil ﬂora, 329–330 Comite´ Spe´cial de l’Anne´e Ge´ophysique Internationale (CSAGI), 23–24

Computer modelling CO2 and ice-sheet inception at E/O boundary, 344–348 Continental rise glacial stratigraphy, Wilkes Land, 132–133 Continental shelf glacial stratigraphy, Wilkes Land, 130–131 Continental slope deposits, 159 Continental slope glacial stratigraphy, Wilkes Land, 131 Contribution to global sea level, isostatic modelling, 558–559 Crary Mountains, 423 Crary Trough Mouth Fan channel-levee complexes, 151 seismic line NARE-8517 across, 148 Cratons, 261 Crozet basin, 102 CRP-2A drill core environmental proxy data, Oligocene–Miocene boundary, 381, 382 revised age models for, 383 CRP-3 drill core, 317–321 Cryosphere, 534 CSAGI (see Comite´ Spe´cial de l’Anne´e Ge´ophysique Internationale (CSAGI)) Cytadela leaf ﬂora, 328–329 Dansgaard-Oeschger (D-O) interstadials, 467 dD (deuterium) signal, 552 Deep ice cores proxy climate records from, 53 Deep Sea Drilling Project (DSDP), 40–41 Wilkes Land margin, 133 Deep-sea isotopes Antarctic Ice Sheet proxy records, 50

Subject Index

Deep western boundary currents (DWBC), 101–103, 490 Depositional model for Late Oligocene/Early Miocene, 319–320 Depositional patterns, 159 of Antarctic Ice Sheets, 49 Dinocysts Late Eocene–Early Oligocene, 337–339 Middle Eocene, 339–340 Palaeocene–Middle Eocene, 336–337 D-O climate oscillations, 467 D-O interstadials (see Dansgaard-Oeschger (D-O) interstadials) Dragon Glacier Moraine, macroﬂora deposits, 328 Drake Passage, climate evolution, 346 Drake Plate, 275 Drill cores, in western Ross Sea, 317–321 Drilling climate records, 316–321 in Prydz Bay, 321–325 records of Ross Sea ice-sheet evolution, 122–126 on Wilkes Land margin, 133 Dronning Maud Land continental slope, 145 Dronning Maud Land margin, 152 seismic line AWI-90110 across, 147 Dry Valley Drilling Project (DVDP), 43, 469 DSDP (see Deep Sea Drilling Project (DSDP)) DSDP Leg 28, 40, 43 DSDP Leg 29, 40

581

DSDP–ODP drilling dates initiation of Antarctic Ice Sheet, 50–52 DVDP (see Dry Valley Drilling Project (DVDP)) DVDP-15, 45 DWBC (see Deep western boundary currents (DWBC)) EAIS (see East Antarctica ice sheet (EAIS)) Early glacial deposites during late Eocene, Prydz Bay, 138–139 ice-sheet evolution in Ross Sea region during late Oligocene to early Miocene, 123–124 Earth Models of Intermediate Complexity (EMIC), 505 Earth’s temperature, variation from cenozoic to present, 3 East Antarctica, 259–260 East Antarctic Craton, 261 East Antarctic Ice Sheet (EAIS), 403–404, 472, 573 marine record, Middle Miocene to Pliocene, 424–428 East terrestrial environments, Middle Miocene to Pliocene, 404–417 East wind drift, 100 EBM (see Energy-balance climate models (EBM)) Edward VII Peninsula (Westernmost Marie Byrd Land), 271 Eemian interglacial, 536–537 EISMINT (see European Ice Sheet Modelling Initiative (EISMINT)) ELA (see Equilibrium line altitude (ELA)) Ellsworth Land, 421–422 Ellsworth Orogeny, 274–275 Ellsworth–Pensacola Mountains, 274

582

Subject Index

EMIC (see Earth Models of Intermediate Complexity (EMIC)) Enderby land, 537–538, F12.2 Energy-balance climate models (EBM), 501 Environmental records from marine microfossils, 335–342 from palynomorphs, 331–335 Eocene environment in Ross Sea, 314–316 ocean isotope record, climate evolution, 342–344 Eocene erratic rocks found in coastal region of Ross Sea, 122–123 Eocene/Oligocene climate signals from sedimentary record, 311–325 EPICA (see European Programme for Ice Coring in Antarctica (EPICA)) Equilibrium line altitude (ELA), 250–251 Erebus volcanic province, Ross Sea tectonic evolution, 283–284 Erratic rocks, Eocene found in coastal region of Ross Sea, 122–123 Eucampia antarctica, 472 European Ice Sheet Modelling Initiative (EISMINT), 242–245 outcomes, 244–245 European Programme for Ice Coring in Antarctica (EPICA), 469 Eustatic sea-level curves, Oligocene–Miocene boundary, 378–379 Exhumation events, variation Transantarctic Mountains, 286

Ferrar Supergroup, 277 Filchner-Ronne Ice Shelf, 532, 548–549, 550, F12.1 Flow of ice, 236 in central regions, 245–247 processes controlling, 237 radar layering and internal, 247 Fluctuation, NH ice volume, 468 Fossil Hill Formation, 329 Fossil plants climate records, Eocene/Oligocene, 325–331 Fossil record, terrestrial of east Antarctic climate, 414–417 Framnes mountains, 538 Franz Joseph Land, 15 Gaussberg, 541 GCM (see Global Climate Model (GCM)) Geological information, 535–550 Holocene glacier, climate ﬂuctuations, 549 Last Glacial Maximum, 537–549 last interglacial, 536–537 late Quaternary ice-sheet, 535–536 pattern and timing, of ice retreat, 550 Prydz Bay, 550 Geological map, 262 Geological time periods, Cenozoic, 2 Geomagnetic Polarity Time Scale calibration history, Oligocene– Miocene boundary, 373 Geotectonic settings before Gondwana amalgamation, 261–264 present-day, 259–260 George V Land, 541 Glacial climate, 533–534 Glacial history of Antarctic, 4–9

Subject Index

Glacial-interglacial cycles, 488–499 (see also Oceanic variability, from southern ocean sediment cores) iceberg discharge and migration, 498–499 intermediate waters, 491 ocean fronts, 496–497 surface ocean currents, 497 surface waters, 492–495 thermohaline circulation, 488–490 Glacial-interglacial (G-I) cycles, 468 Glacial-interglacial time scales, 481–485 Glacial sedimentation ancestral Lambert Glacier, advance and recession stages, 323–324 Glacial sequences on Continental shelf, 130–131 Glacial stratigraphy continental rise, 132–133 continental shelf, 130–131 continental slope, 131 Glacial variability, from continental margin geological record, 469–479 ANDRILL, 476–479 Antarctic Peninsula, 472–473 Prydz Bay, 473–475 Ross Sea, 475–476 Weddell Sea, 471–472 Glaciation record of Wilkes Land, 134–135 GLIMMER community ice-sheet model, 555 Global Circulation Model (GCM) simulation, 347, 348 south polar climate variations in response to a growing ice sheet, 388 Global Climate Model (GCM), 504, 551

583

Global ice volume evolution during Eocene to Oligocene, 342–344 Global positioning system (GPS), 559 GLOCHANT (see Group of Specialists on Global Change (GLOCHANT)) Glomar Challenger, 40, 42 Gondwana amalgamation, 270 geotectonic settings before, 261–264 break-up, 276–287 tectonic stages of, 276–279 dispersal of Southern Hemisphere Continents, 276–287 supercontinent, 265–276 Gondwana, evolution late precambrian–early palaeozoic (c. 600–450Ma), 266–272 from Rodinia, 267–272 upper palaeozoic–mesozoic (c. 450–180Ma), 272–276 GPS (see Global positioning system (GPS)) Grenvillian orogen, 263 Grounding-line fan model of glaciomarine sedimentation, Late Oligocene/Early Miocene, 319–320 Group of Specialists on Global Change (GLOCHANT), 54 Gyres, 99–100 ‘‘Heroic era’’, 35, 38 Holocene, 531 glacier, climate ﬂuctuations, 549 Iceberg discharge and migration, 498–499 (see also Glacial-interglacial cycles) Iceberg-rafted debris (IRD), 473

584

Subject Index

Ice cores proxy climate records from deep, 53 Ice coring, future challenges, 487–488 Ice-sheet development mid-early Miocene to early Pliocene in Ross Sea region, 124–125 Ice-sheet evolution in Ross Sea early glacial, 123–124 during mid-early miocene to early pliocene, 124–125 polar ice sheet, 125–126 pre-ice-sheet, 122–123 Ice sheet expansion, 509 (see also Antarctic climate evolution) Ice-sheet ﬂuctuations Late Miocene–Early Pliocene, 427–428 Ice-sheet hysteresis inﬂuence across Oligocene–Miocene boundary, 390–391 Ice-sheet inception at E/O boundary, computer modelling, 344–348 Ice sheet Model Intercomparison Programme (ISMIP), 245 Ice-sheet modelling, 55–56, 550–551 Ice-sheet models, numerical, 238–241 versus Antarctic glaciological data, 245–249 Ice-sheet processes, 236–238 Ice Sheets, Antarctic East (see East Antarctic Ice Sheet (EAIS)) modelling, 437–441 West (see West Antarctic Ice Sheet (WAIS)) Ice-sheet sensitivity, 554 Ice streams, 154

Ice volume, global evolution during Eocene to Oligocene, 342–344 Ice-volume variations, 170–172 ICSU (see International Council of Scientiﬁc Unions (ICSU)) IGY (see International Geophysical Year (IGY)) IMAGES programme (see International Marine Global Change Study (IMAGES) programme) IMO (see International Meteorological Organization (IMO)) Indian DWBC inﬂow, 102 Indo-Antarctica, 270 InSAR (see Interferometric Synthetic Aperture Radar (InSAR)) Integrated Ocean Drilling Programme (IODP), 65, 469 Interferometric Synthetic Aperture Radar (InSAR) for surface velocity of ice sheets, 248–249 Interglacial ice sheet, 551–552 International Council of Scientiﬁc Unions (ICSU), 23, 25 International Geophysical Year (IGY), 1957–1958, 22–26 establishment of SCAR, 25 estimation of ice volume in Antarctica, 25 International Marine Global Change Study (IMAGES) programme, 8 International Meteorological Conference, 1879, 17 International Meteorological Organization (IMO), 19 International Polar Commission, 17, 19–20

Subject Index

International Polar Conference, 17 International Polar Year (IPY) 1882–1883, 14–19 1932–1933, 19–22 1957–1958, 22–26 2007–2008, 27–30 themes, 29 IODP (see Integrated Ocean Drilling Programme (IODP)) Isostasy, 238 Isostatic modelling contribution to global sea level, 558–559 of post-glacial uplift, 558–559 Isotopic record Oligocene–Miocene boundary, 374–377 James Ross Island Volcanic Group, 417, 421 John Mercer’s hypothesis for behaviour of Antarctic Ice Sheet, 47 Jurassic Ferrar Large Igneous Province (FLIP), 277 Jutul Penck Graben, 279 K–Ar dating for Antarctic Cenozoic ice sheets, 39 King George Island fossil plants, 328–330 stratigraphy, 313–314 Krakowiak Glacial Member, 314 Lake glubokoye, 537 Lambert Glacier, ancestral glacial sedimentation, advance and recession stages, 323–324 Lambert glacier-Amery ice shelf, 7, 538–539

585

Lambert Glacier region Middle Miocene to Pliocene history, 410–414 Lambert Graben, 278–279 La Meseta Formation, 312 Land-based drilling to study Antarctic Cenozoic glaciation, 43 Landscape modelling of Antarctic Ice Sheet, 58–59 Larsemann Hills, 539, 549–550 Larsen Shelf sediments, 149 seismic line AWI-97051 across, 147 Last glacial maximum (LGM), 8–9, 473, 531, 537–549, 574 Antarctic Peninsula, 545–547 Coats Land, 548–549 Enderby Land, 537–538, F12.2 Filchner-Ronne Ice Shelf, 548–549 Lambert Glacier-Amery Ice Shelf, 538–539 Mac.Robertson Land, 538–539 Marie Byrd Land, 544–545 Princess Elizabeth Land, 539–542 Prydz Bay, 538–539, F12.3 Queen Mary Land, 539–542 Queen Maud Land, 537–538, F12.2 in Ross Sea, 125–126 Ross Sea Embayment, 542–544 Transantarctic Mountains, 542–544 Weddell Sea Embayment, 548–549 Wilkes Land, 539–542 Late Eocene–Early Oligocene dinocysts, 337–339 Late Eocene-early Oligocene cooling, 5 Late Miocene–Early Pliocene ice-sheet ﬂuctuations, 427–428 marine records, 435–436

586

Subject Index

Late Oligocene/Early Miocene depositional model for, 319–320 Victoria Land coast in, 321 Late Palaeocene-Late Eocene fossil ﬂoras, Antarctic Peninsula region, 326–327 Late Pliocene cooling, 509 (see also Antarctic climate evolution) Late Pliocene-early Pleistocene G-I climate cycles, 509–510 (see also Antarctic climate evolution) Late Precambrian–Early Palaeozoic (c. 600–450 Ma) Gondwana, evolution, 266–272 palaeo-paciﬁc margin, 270–272 Late Quaternary ice-sheet, 535–536 reconstructions, 533–534 Law dome, 105, 534 LCDW (see Lower circumpolar deep water (LCDW)) LGM (see Last glacial maximum (LGM)) Lithostratigraphic changes, in antarctic margin, 169 in drill cores from continental rise, 167–168 facies, 170 sea-level and ice-volume changes, 170–172 ‘‘Longyear 44’’ drilling rig, 43, 46 Lower circumpolar deep water (LCDW), 94 Lu¨tzow Holm–Prydz Bay Pan-African orogenic belt, 268–270 Mac.Robertson Land, 538–539 Magmatism in Ross Sea Sector, during the Cenozoic, 279–284

Magnesium/Calcium (Mg/Ca) data across Oligocene–Miocene boundary, 375–377, 380 Marie Byrd Land, 422–424, 544–545 Marine ice sheet margin development, 509 (see also Antarctic climate evolution) Marine Isotope Stage 31, 510–511 (see also Antarctic climate evolution) Marine microfossils climate records from, 335–342 Marine record, Middle Miocene to Pliocene Antartic Peninsula, 429–433 EAIS, 424–428 Southern Ocean, 433–436 WAIS, 429–433 Marine records, Late Miocene–Early Pliocene, 435–436 Mass balance, 237–238 Maud Rise, records from, 380 MCMC (see Monte Carlo Markov Chain (MCMC)) McMurdo Dry Valleys, 43 McMurdo Erratics, 315 McMurdo Ice Shelf (MIS), 471 McMurdo Sound climate data, Oligocene–Miocene boundary, 381–385 fossil ﬂora, 330–331 map and cross-section of, 44 post-IGY surveys of glacial deposits in, 39 stratigraphy, 284 McMurdo Sound Sediment and Tectonic Studies Project (MSSTS), 46–47 MCS (see Multichannel seismic (MCS) reﬂection) MCS proﬁles, location of Wilkes Land margin, 128

Subject Index

Mertz glaciers, 541 Mesozoic–Cenozoic sedimentary basin, 152 Methane concentrations, 483 Meyer Desert Formation, fossil records, 416 Microfossils, marine, climate records from, 335–342 Middle-late Miocene cooling, 6 Middle Miocene Climate Transition (MMCT), 433–435 Middle Miocene cooling, marine records, 425–427 Middle Miocene to Pliocene terrestrial environments, 404–424 east, 404–417 west, 417–424 Mi-1 d18O event, origin, 386 Mid-Pleistocene Transition (MPT), 468, 505, 511–512 (see also Antarctic climate evolution; Pleistocene ice volume variation modelling) Mi-1 glaciations (see Miocene (Mi-1) glaciations) Millennial-scale climate variability, 485–486, 506, 512–513 (see also Antarctic climate evolution; Pleistocene ice volume variation modelling) Miocene climatic optimimum, 62–63 Miocene (Mi-1) glaciations, 5–6 Modelling Antartic climate and ice sheets, 437–441 CO2 and ice-sheet inception at E/O boundary, 344–348 Monte Carlo Markov Chain (MCMC), 502 Moraine complexes, 152

587

Mount Murphy, 422–423 Mount Wawel, macroﬂora deposits, 328 Moutone’e Lake, 549 Mozambique suture zone, 267–268 MPT (see Mid-Pleistocene Transition (MPT)) MSSTS (see McMurdo Sound Sediment and Tectonic Studies Project (MSSTS)) Multichannel seismic (MCS) reﬂection, 117 study of Ross Sea, 119–126 ice-sheet evolution, 122–126 NADW (see North Atlantic Deep Water (NADW)) Neogene age volcanic outcrops in West Antarctica, 418–420 Neogene (15-2 Ma) distribution terrestrial volcanic outcrops, West Antarctica, 413 Neogene stratigraphic section of Ross Sea, 126 Neogloboquadrina pachyderma, 474 New Zealand Antarctic Programme (see McMurdo Sound Sediment and Tectonic Studies Project (MSSTS)) NH (see Northern Hemisphere (NH)) Nimrod Glacier area, 270–271 Ninnis glaciers, 541 North Atlantic Deep Water (NADW), 488 Northern continental slope, 154 Northern Hemisphere (NH), 467 ice volume ﬂuctuation, 468 summer insolation, 467 North Greenland Ice Core Project (North GRIP), 481

588

Subject Index

Novolazarevskaya ice shelf, 537 Nowaja Semlja, 15 Numerical ice-sheet models, 238–241 Numerical modelling of Antarctic Ice Sheet, 235–253 EISMINT, 242–245 ice-sheet models, 238–241 versus glaciological data, 245–249 ice-sheet processes, 236–238 ice-sheet reconstructions, 249–253 model input, 241–242 Numerical modelling reconstructions, 550–559 Antarctica modelling, 555–558 ice-sheet modelling, 550–551 ice-sheet sensitivity, 554 interglacial ice sheet, 551–552 time dependent simulations, 552–554 Ocean circulation, 467 inﬂuence across Oligocene–Miocene boundary, 385–389 Ocean Drilling Programme (ODP), 469 age-depth model for drilling site of, 141 and ANTOSTRAT, 53–54 deep-sea isotope data from sites in Paciﬁc and Atlantic Oceans, 50 drilling sites of Prydz Bay, 136 Oceanic variability, from southern ocean sediment cores, 488–501 glacial-interglacial cycles, 488–499 millennial-scale cycles, 499–501 Ocean isotope record climate evolution during Eocene and Oligocene, 342–344

Ocean temperatures evolution during Eocene to Oligocene, 342–344 ODP (see Ocean Drilling Programme (ODP)) Oligocene glaciation, 158 ice sheets dynamic, 49–50 ocean isotope record, climate evolution, 342–344 Oligocene–Miocene boundary drivers of climate change across, 385–391 identiﬁcation of, 371–374 proxy records, 374–379 records from Antartic margins, 379–385 Oligocene–Miocene boundary Mi-1 glaciation, 5–6 Orbital parameters inﬂuence across Oligocene–Miocene boundary, 389–390 Orogenic belts, 263 Oxygen isotope data across Oligocene–Miocene boundary, 374–376, 380 Oxygen isotope records for Cenozoic, 343 Paciﬁc–Antarctic Ridge, initiation, 282 Paciﬁc DWBC inﬂow, 102–103 Pagodroma Group, 7, 404–405 clay mineralogy, 412–414 glacial erosional features, 411 Middle Miocene to Pliocene history, 410–414 sedimentary characteristics, 411 stratigraphy, 408 Palaeoceanography, 339–342 Palaeocene–Middle Eocene dinocysts, 336–337

Subject Index

Palaeogene, Antarctic strata pollen assemblages, 332–333 Palaeogene fossil plants, 325–328 Palaeo-ice streams and S3/S2 boundary, 159 Palaeo-paciﬁc margin Gondwana evolution, late precambrian-early palaeozoic, 270–272 Palaeo sea level records, Oligocene–Miocene boundary, 377–379 Palynomorphs climate records from, 331–335 from OPD drilling site in Prydz Bay, 139 Pattern and timing, of ice retreat, 550 Payer, Julius, 16 Plant macrofossils climate records from, 325–328 Plate-tectonically active part, 275–276 Pleistocene glacial cycles, 7–8 ice sheets, 251–253 Pleistocene ice volume variation modelling, 501–509 Antarctica, 506–509 millennial-scale variability, 506 MPT, 505 Northern Hemisphere, 501–505 Pliocene ice sheets, 249–251 record, of Antarctic Ice Sheet, 7 vegetation distributions, climate modelling, 439 Plio-pleistocene strata, 543 Point Hennequin Flora, 328 Polar ice sheet evolution in Prydz Bay, 142–143

589

evolution in Ross Sea region during early Pliocene through Quaternary, 125–126 Polar stereographic projection Oligocene–Miocene boundary locations, 372 Pollen assemblages in Antarctic Palaeogene strata, 332–333 Polonez Cove Formation, 314 Post-glacial uplift, isostatic modelling, 558–559 Pre-ice-sheet Cenozoic shelf edge, 148 evolution in Ross Sea region during pre-late-Oligocene, 122–123 palaeo-climate evolution of Prydz Bay during pre-late Eocene, 137–138 stratigraphy of Wilkes Land, 127–130 Prince Gustav Channel, 549 Princess Elizabeth Land, 539–542 Proxy records Oligocene–Miocene boundary, 374–379 Prydz Bay (PB), 473–475, 538–539, F12.3 cenozoic climate history, seismic interpretation and drilling, 135–144 ice-sheet development during Oligocene–Miocene, 139–142 drilling in, 321–325 geomorphic and seismic stratigraphic features of, 163 geomorphic features of, 163, 165 lithostratigraphic transitions in, 168–169 sea-level variations, 171

590

Subject Index

seismic lines and ODP drill sites of, 136 seismic stratigraphic features of, 165–167 Quaternary Antarctic evolution, 553–554 Quaternary climate dataset, 468 Queen Mary Land, 539–542 Queen Maud Land, 537–538 Radar layering Antarctic Ice Sheet, 247 RCM (see Regional Climate Model (RCM)) ‘‘Refrigerator trap’’, 342 Regional Climate Model (RCM), 504 Rennick Graben, 279 Rodinia Gondwana, evolution from, 267–272 Ross embayment McMurdo Sound coastal area of, 284 rifting in, 280–282 Ross gyre, 100 Ross Ice shelf ANDRILL drilling on, 124 Ross Ice Shelf Project (RISP), 46 Ross Orogeny, 271 Ross Sea, 475–476, 532, 542–544, F12.1 Cenozoic climate history, seismic interpretation and drilling data, 118–126 continental shelf edge, 166 Eocene environments in, 314–316 fossil ﬂora from McMurdo Sound, 330–331 geomorphic and seismic stratigraphic features of, 165

lithostratigraphic transitions in, 168–170 magmatism, during Cenozoic, 279–284 sea-level variations in, 171 seismic data used in glacial sedimentary processes, 121–122 ship-based drilling at eastern, 40 stratigraphic correlation, 121 structural framework of, 119 tectonic evolution, during Cenozoic, 279–284 Ross Sea, South Western climate data, Oligocene–Miocene boundary, 381–385 Ross Sea, western drill cores in, 317–321 Ross Sea Rift, 279 RSL curve, 549 Salinity trends, of Southern Ocean, 105–106 SAM (see Southern Annular Mode (SAM)) SAMW (see Subantarctic mode water (SAMW)) Satellite interferometry to calculate ice velocity, 248–249 SCAR (see Scientiﬁc Committee on Antarctic Research (SCAR)) Schirmacher oasis, 537 Scientiﬁc Committee on Antarctic Research (SCAR) CIROS, 47–48 establishment during IGY, 25 planning for future research, 52–53 Sea ice drilling, 54–55 future prospective of, 64

Subject Index

to study Antarctic Cenozoic glaciation, 43–45 advancement during twenty-ﬁrst century, 59–63 Sea-level variations, 170–172 Sea-riser technology, 471 Sedimentary record (Eocene/Oligocene) climate records from, 311–325 Seismic stratigraphy Antarctic Ice Sheet proxy records, 50 ice-sheet evolution in Ross Sea region, 122–126 sea level changes, 46–47 Seymour Island fossil ﬂora, 330 Shelf aggradation, 145, 147 Shelf-ice drilling future prospective, 64 to study history of Antarctic Cenozoic Glaciation, 59–63 Ship-based drilling CIROS Project, 47–48 future prospective of, 64–66 to study Antarctic Cenozoic glaciation, 40–43 Siliceous biogenic facies, 149 Single-channel seismic (SCS) surveys, of Ross sea, 119 Siple coast, of Ross Sea, 537 Sirius Group, 7, 404–405 glacial erosional features, 407 Middle Miocene to Pliocene history, 406–410 sedimentary characteristics, 407 stratigraphy, 408 Skelton Glacier area, 270 Sørsdal Glacier, 541 Southern Annular Mode (SAM), 104 Southern Ocean circulation, 97–103 Antarctic circumpolar current, 97–99

591

subpolar circulation, 99–100 thermohaline circulation, 100–103 climate change/warming, 103–107 oceanographic elements of, 87 salinity trends of, 105–106 water mass formation and dispersal, 88–97 South Orkney Islands region, 155–156 South polar seasonal temperatures, sea ice and winds in response to a growing ice sheet, GCM simulation, 388 South Shetland Islands, 275 stratigraphy, 313–314 South Shetland Islands region Bransﬁeld Strait, 156–157 scientiﬁc drilling in, 158 South Western Ross Sea climate data, Oligocene–Miocene Boundary, 381–385 SPECMAP reconstruction, 553 Sporomorphs, in Tasman Sea relative abundance, 335 S3/S2 boundary, 159 Stratigraphic correlation of Antarctic sedimentary strata with other Gondwana continents, 38–39 of Ross Sea area, 121 Stratigraphy continental shelf glacial, Wilkes Land, 130–131 Wilkes Land Acoustic, 127–133 Stratigraphy, seismic Antarctic Ice Sheet proxy records, 50 Strike–slip fault system in Victoria and Oates Lands, 279 Structural framework of Ross Sea, 119

592

Subject Index

Subantarctic mode water (SAMW), 90–92, 491 Subglacial lakes location of, 247–248 sediments deposites, 63 Subpolar circulation, 99–100 Summer insolation, NH, 467 Supercontinent, Gondwana, 265–276 Surface ocean currents, 497 (see also Glacial-interglacial cycles) Tasmanian Gateway climate evolution, 346 Tasman Sea sporomorphs, relative abundance, 335 Taylor Dome, 534 Taylor Dry Valley, 43 Tectonic evolution in Ross Sea Sector, during Cenozoic, 279–284 of Transantarctic Mountains, 284–287 Tectonic isolation inﬂuence across Oligocene–Miocene boundary, 385–389 Tectonic stages of Gondwana break-up, 276–279 Terrestrial environments, Middle Miocene to Pliocene east, 404–417 west, 417–424 Terrestrial realm climate records from, 325–335 Terrestrial volcanic outcrops in West Antarctica distribution of Neogene (15-2 Ma), 413 Tethyan closure, 389

Thermohaline circulation (THC), 100–103, 488–490 (see also Glacial-interglacial cycles) Time dependent simulations, 552–554 Transantarctic ﬂora, 336 Transantarctic mountains, 259–260, 542–544 exhumation events, variation, 286 map of, 281 Middle Miocene to Pliocene history, 406–410 new records from, 63 reinterpretation of climate discoveries in, 56–57 tectonic evolution of, 284–287 wet-based glacial deposits, 48–49 UCDW (see Upper circumpolar deep water (UCDW)) Upper circumpolar deep water (UCDW), 94 Upper Palaeozoic–Mesozoic (c. 450–180 Ma) Gondwana, evolution, 272–276 Vegetation distributions Pliocene, climate modelling, 439 Velocity map of Antarctic Ice Sheet, 68 Vestfold Hills, 542, 549 Victoria Group, 274 Victoria Land coast in Late Oligocene/Early Miocene, 321 Volcanic outcrops of Neogene age in West Antarctica, 418–420 Vostok core, 53 Vostok ice core, 551–552

Subject Index

WAIS (see West Antarctic Ice Sheet (WAIS)) Water mass, formation and dispersal, 88–97 Antarctic bottom water, 94–97 circumpolar deep water, 92–94 Subantarctic mode water and Antarctic intermediate water, 90–92 surface water, 88–90 Weddell–Enderby Basin, 101 Weddell gyre, 99 Weddell Orogeny, 274–275 Weddell Sea, 471–472 Cenozoic climate history, seismic interpretation and drilling data, 144–152 continental margin sediments, 151 environmental change in Miocene sediments, 151 Oligocene sediments, 149 sedimentation rates, 150 glacigenic sediments, acoustic stratiﬁcation of wedge of Crary Trough, 145, 148 Dronning Maud Land margin and Larsen Shelf, 147 ice-sheet history, 151–152 MCS lines, 145 pre-ice-sheet Cenozoic shelf edge, 148 pre-Oligocene section of inferred turbidites, 149 Weddell Sea Embayment, 548–549 West Antarctica, 259–260 West Antarctica, terrestrial volcanic outcrops distribution of Neogene (15-2 Ma), 413

593

West Antarctic Ice Sheet (WAIS), 118, 404, 536, 572 marine record, Middle Miocene to Pliocene, 429–433 West Antarctic Rift System (WARS), 259 map of, 281 Westernmost Marie Byrd Land (Edward VII Peninsula), 271 Western Ross Sea drill cores in, 317–321 West terrestrial environments, Middle Miocene to Pliocene, 417–424 West wind drift, 97 Weyprecht, Karl, 14–17 Whitmore Mountains, 274 Wilkes Land, 539–542 acoustic stratigraphy, 127–133 Cenozoic climate history, seismic interpretation and drilling data, 127–135 drilling on margins of, 133 geomorphic and seismic stratigraphic features of, 163, 165 inferred long-term record of glaciations, 134–135 seismic stratigraphic history, 166 Wilkes Land margin drilling on, 133 location of MCS proﬁles, 128 Wilkins ice shelf, 572, F13.2 Wind drift, 97 Windmill Islands, 541 WMO (see World Meteorological Organization (WMO)) World Meteorological Organization (WMO), 19, 27

Developments in Earth & Environmental Sciences, 8 F. Florindo and M. Siegert (Editors) r 2009 Elsevier B.V. All rights reserved DOI 10.1016/S1571-9197(08)00013-X

Chapter 13

Concluding Remarks: Recent Changes in Antarctica and Future Research Fabio Florindo1, and Martin Siegert2 1

Istituto Nazionale di Geoﬁsica e Vulcanologia, via di Vigna Murata 605, 00143 Roma, Italy 2 School of GeoSciences, Grant Institute, University of Edinburgh, The King’s Buildings, West Mains Road, Edinburgh EH9 3JW, UK

ABSTRACT Antarctic Climate Evolution has been, and will be, hugely inﬂuential in the development of Earth’s environment. This book has detailed how Antarctica changed during several key stages in the Cenozoic. Here we take stock of past changes and consider how they may be helpful in evaluating future changes in Antarctica.

Warming of the climate system is unequivocal, as is now evident from observations of increases in global average air and ocean temperatures, widespread melting of snow and ice in Greenland and in the Antarctic Peninsula, and rising global sea levels. Eleven of the 12 years between 1995 and 2006 are among the 12 warmest years in the instrumental record since 1850 (IPCC, 2007). According to the latest IPCC worst-case scenario projections (i.e. continued greenhouse gas emissions at or above current rates), global annual mean temperatures by 2100 are likely to exceed those that have been experienced by the Earth in the last 40 myr, that is before the Antarctic Ice Sheet ﬁrst developed (IPCC, 2007). If warming continues, the implication is that ice loss to the ocean may far outweigh gain from ice accumulation, with huge implications for global sea level (and associated Corresponding author. Tel.: þ39 0651860 383; Fax: þ39 0651860 397;

E-mail: ﬂ[email protected] (F. Florindo).

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Figure 13.1: Ice-sheet surface elevation changes, recorded by satellite altimetric measurements, between 1992 and 2003. Note how ice loss is concentrated at the margins of ice streams in the Amundsen Bay region of East Antarctica, and the Totten and Cook glaciers in East Antarctica. These regions of ice loss are characterised by ice resting on a bed suppressed several hundred metres below sea-level, implying an ice-ocean connection to the current retreat (from Shepherd and Wingham (2007), with permission of the authors). major changes in coastlines and the inundation of low-lying areas), atmospheric composition and dynamics, and ocean circulation. Recent comprehensive, and near continuous, satellite altimetric observations indicate that the Antarctic Ice Sheet is losing mass, and that the rate of loss has increased steadily during the last few decades (Fig. 13.1, e.g. Davis et al., 2005; Shepherd and Wingham, 2007; Rignot et al., 2008). Most loss is from the West Antarctic Ice Sheet (WAIS), especially along the Bellingshausen and Amundsen seas (132760 Gt yr1 in 2006). During the past ca. 40 years many glaciers, and a series of ice-shelves, have retreated in the Antarctic Peninsula (e.g. Cook et al., 2005) and, since the new millennium, two large ice shelves, located either side of the Antarctic Peninsula (the Larsen B and Wilkins ice shelves, with extents of 3,250 km2 and over 570 km2, respectively; Fig. 13.2), have suddenly broken up in dramatic fashion and separated from the continent. The decay of these ice shelves will not have affected sea level because ice shelves are aﬂoat and thus already displacing their weight of water. However, ice-shelf decay has the potential to greatly reduce buttressing forces on grounded ice that resist the ﬂow from the continent to the ocean. In other words, ice-shelf decay may encourage further loss of grounded ice, which will affect global sea level. A number of factors contribute to ice-shelf collapse, including atmospheric and oceanic

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Figure 13.2: High-resolution, enhanced-colour image of the Wilkins Ice Shelf in Antarctica, taken on 8 March 2008 by Taiwan’s Formosat-2 satellite (Credit: left, National Snow and Ice Data Center; right, National Snow and Ice Data Center/Courtesy Cheng-Chien Liu, National Cheng Kung University (NCKU), Taiwan and Taiwan’s National Space Organization (NSPO); processed at Earth Dynamic System Research Center at NCKU, Taiwan).

temperatures. If these continue to rise, we may see more ice shelves decaying across the Antarctic Peninsula and elsewhere in Antarctica (e.g. Morris and Vaughan, 2003; MacAyeal and others, 2006; Glasser and Scambos, 2008). It is worth noting that, over the past 50 years, the western part of the Antarctic Peninsula has experienced the greatest temperature increase on Earth, rising by nearly 31C (e.g. King et al., 2003; Turner et al., 2005). This is approximately 10 times the mean rate of global warming, as reported by the IPCC. The temperature increase is not conﬁned to the Antarctic Peninsula; it is also occurring in and over the oceans surrounding Antarctica. For example, it is now well-established that the waters of the Antarctic Circumpolar Current (ACC) are warming more rapidly than the global ocean as a whole. A comparison of temperature measurements from the 1990s with data from earlier decades shows a large-scale warming of around 0.21C in the ACC waters at around 700–1,100 m depth (Gille, 2002). Across the East Antarctica Ice Sheet (EAIS) the mass balance is near a steady state with the ﬂow of ice. Its interior is increasing in overall thickness, which is probably the result of increased ice accumulation due to warmer (and moister) atmospheric conditions over the continental interior (Church et al., 2001). This mass gain is balanced by ice thinning/loss across the potentially dynamic marine sectors in Wilkes Land (Davis et al., 2005). Further atmospheric and ocean warming may be incompatible with

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sustaining mass balance in this region, implying it too is susceptible to further changes in a warming world. A greater understanding of past changes in this region of East Antarctica is clearly warranted. Comprehending glacial history since the Last Glacial Maximum (LGM), which culminated around 20,000 years ago, is also critical to knowing whether recent changes are a result of human activities and/or a result of icesheet relaxation since the LGM. The geological record tells us that the WAIS retreat occurred during the Holocene, far later than Northern Hemisphere ice sheets. It is certainly possible that changes observed in some locations today can be at least part explained by post-LGM ice retreat to a more stable interglacial condition. Geological evidence on the rate of ice loss since the LGM is therefore needed at several sites across the Antarctic continent. Understanding Antarctic climate evolution requires continent-wide as well as continent-to-deep-sea studies of past climate records, to decipher their separate but related histories. Some of these activities are long-term undertakings but commitment within the frame of IPY is essential to laying the groundwork for long-term success. Over the next few years ACE aims to integrate geoscience data provided by sediment cores (ANDRILL, andrill.org; SHALDRILL, www.shaldril.rice.edu; IODP, www.iodp.org) and ice cores (EPICA, www.esf.org; ITASE, www2.umaine.edu/itase/; TALDICE; WAIS Divide, www.waisdivide.unh.edu/) with the new generation of climate models that couple atmosphere, ocean, ice sheet and sediment modelling on key time periods from the distant past (i.e. tens of millions of years ago when global temperature was several degrees warmer than today) to the recent past (i.e. during the Holocene, prior to anthropogenic impacts as well as at the LGM). Acquisition of sedimentary records from the ﬂoors of subglacial lakes is also anticipated in the next few years. Such information may inform us about ice-sheet histories, and when West Antarctica was last deglaciated. Drilling, sampling and studying subglacial lakes remotely and without causing their contamination represent a considerable technological challenge. In this regard the exploration of sub-ice lakes represents a good analogue for the exploration of planets and satellites such as Europa. A number of subglacial lakes have been identiﬁed in Antarctica and, of these, one subglacial lake in West Antarctica, named Lake Ellsworth (from the American explorer Lincoln Ellsworth), is well suited to exploratory research (www.geos.ed.ac.uk/ellsworth). This book has demonstrated that the Antarctic Ice Sheet has undergone many changes, and has varied in size considerably over the Cenozoic. The nature of these changes has been shown to be associated with global climate conditions, forced externally or through interaction with ice and climate

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conditions in Antarctica and the Southern Ocean. IPCC predictions of atmospheric greenhouse gases are pertinent to future ice volumes in Antarctica. If such gases continue to rise, in a few centuries time the value of atmospheric carbon dioxide may be greater than at any time in the Cenozoic. The obvious danger is that, in a warming world, the Antarctic Ice Sheet may respond to climate and ocean changes as it has done in the past. In other words, the palaeo-ice-sheet reconstructions highlighted in this book may be relevant to assessments of future changes in Antarctica. Over the next decade, ACE will be pursuing a broad range of objectives to better comprehend past Antarctic changes, through the organisation of workshops, where interdisciplinary research can be discussed, and through publications (e.g. Florindo et al., 2003, 2005, 2008; Barrett et al., 2006), allowing dissemination of results to a wide audience. It is only through such integration of geological data and numerical modelling that quantitative assessments of past changes, and possible future scenarios, can be achieved.

REFERENCES Barrett, P. J., Florindo, F., & Cooper, A. K. (Eds) (2006). Cenozoic paleoenvironments: Geologic record and models. Palaeogeogr. Palaeoclimatol. Palaeoecol., 231, 1–252. Church, J. A., Gregory, J. M., Huybrechts, P., Kuhn, M., Lambeck, K., Nhuan, M. T., Qin, D., Woodworth, P. L., Anisimov, O. A., Bryan, F. O., Cazenave, A., Dixon, K. W., Fitzharris, B. B., Flato, G. M., Ganopolski, A., Gornitz, V., Lowe, J. A., Noda, A., Oberhuber, J. M., O’Farrell, S. P., Ohmura, A., Opperheimer, M., Peltier, W. R., Raper, S. C. B., Ritz, C., Russell, G. L., Schlosser, E., Shum, C.K., Stocker, T. F., Stouffer, R. J., Van de Wal, R. S. W., Voss, R., Wiebe, E. C., Wild, M., Wingham, D. J., Zwally, H. J. (2001). Changes in Sea level. In: J. T. Houghton, et al (Eds). Climate Change 2001: The Scientiﬁc basis. Cambridge University Press, New York, pp. 639–694. Cook, A. J., Fox, A. J., Vaughan, D. G., & Ferrigno, J. G. (2005). Retreating glacier fronts on the Antarctic Peninsula over the past half-century. Science, 308, 541–544. Davis, C. H., Li, Y., McConnell, J. R., Frey, M. M., & Hanna, E. (2005). Snowfall driven growth in East Antarctic Ice Sheet Mitigates recent sea-level rise. Science, 308, 1898–1901. Florindo, F., Cooper, A. K., & O’Brien, P. E. (Eds) (2003). Antarctic Cenozoic palaeoenvironments: Geologic record and models. Palaeogeogr. Palaeoclimatol. Palaeoecol., 198(1–2), 1–278. Florindo, F., Harwood, D. M., & Wilson, G. S. (Eds) (2005). Long term changes in southern high-latitude ice sheets and climate, the Cenozoic history. Glob. Planet. Change, 45(1–3), 1–260.

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