The Climate of Past Interglacials, Volume 7 (Developments in Quaternary Science)

DEVELOPMENTS IN QUATERNARY SCIENCE, 7 SERIES EDITOR: JAAP J .M . VAN DER MEER THE CLIMATE OF PAST INTERGLACIALS Deve...

Author: F. Sirocko | M. Claussen | T. Litt | M.F. Sanchez-Goni

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DEVELOPMENTS IN QUATERNARY SCIENCE, 7 SERIES EDITOR: JAAP J .M . VAN DER MEER

THE CLIMATE OF PAST INTERGLACIALS

Developments in Quaternary Science (Series editor: Jaap J.M. van der Meer) Volumes in this series 1. The Quaternary Period in the United States Edited by A.R. Gillespie, S.C. Porter, B.F. Atwater, 0-444-51470-8 (hardbound); 0-444-51471-6 (paperback) – 2003 2. Quaternary Glaciations – Extent and Chronology Edited by J. Ehlers, P.L. Gibbard Part I: Europe ISBN 0-444-51462-7 (hardbound) – 2004 Part II: North America ISBN 0-444-51592-5 (hardbound) – 2004 Part III: South America, Asia, Australasia, Antarctica ISBN 0-444-51593-3 (hardbound) – 2004 3. Ice Age Southern Andes – A Chronicle of Paleoecological Events Authored by C.J. Heusser 0-444-51478-3 (hardbound) – 2003 4. Spitsbergen Push Moraines – Including a translation of K. Gripp: Glaciologische und geologische Ergebnisse der Hamburgischen Spitzbergen-Expedition 1927 Edited by J.J.M. van der Meer 0-444-51544-5 (hardbound) – 2004 5. Iceland – Modern Processes and Past Environments ´ . Knudsen Edited by C. Caseldine, A. Russell, J. HarDardo´ttir, O 0-444-50652-7 (hardbound) – 2005 6. Glaciotectonism Authored by J.S. Aber, A. Ber 0-444-52943-8 (hardbound) – 2006 7. The Climate of Past Interglacials Edited by F. Sirocko, M. Claussen, M.F. Sańchez Gon˜i, T. Litt 0-444-52955-1 (hardbound) – 2007 For further information as well as other related products, please visit the Elsevier homepage (http://www.elsevier.com)

Developments in Quaternary Science, 7 Series editor: Jaap J.M. van der Meer

THE CLIMATE OF PAST INTERGLACIALS edited by

Frank Sirocko Institute for Geoscience, University of Mainz, Mainz, Germany

Martin Claussen Meteorological Institute, University Hamburg, and Max Planck Institute for Meteorology, Hamburg, Germany

Marıá Fernanda Sańchez Gon˜i Department of Geology and Oceanography, EPHE-UMR-CNRS 5805, EPOC, University Bordeaux 1, Talence, France

Thomas Litt Institute for Paleontology, University of Bonn, Bonn, Germany

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

Elsevier Radarweg 29, PO Box 211, 1000 AE Amsterdam, The Netherlands The Boulevard, Langford Lane, Kidlington, Oxford OX5 1GB, UK First edition 2007 Copyright ª 2007 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 ðþ44Þ (0) 1865 843830; fax ðþ44Þ (0) 1865 853333; email: [email protected]. Alternatively you can submit your request online by visiting the Elsevier web site at http://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 Library of Congress Cataloging-in-Publication Data A catalog record for this book is available from the Library of Congress British Library Cataloguing in Publication Data A catalogue record for this book is available from the British Library ISBN-13: 978-0-444-52955-8 ISBN-10: 0-444-52955-1 ISSN: 1571-0866 For information on all Elsevier publications visit our website at books.elsevier.com Printed and bound in Italy. 07 08 09 10 11 10 9 8

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Table of Contents Preface Thorsten Kiefer and Christoph Kull ..................................................................... xi Acknowledgements ..........................................................................................................xv Section 1 Forcing Mechanisms (ed. Martin Claussen) 1. Introduction to Climate Forcing and Climate Feedbacks....................................................3 Martin Claussen 2. Insolation During Interglacial ................................................................................................13 A. Berger, M.F. Loutre, F. Kaspar and S.J. Lorenz 3. A Survey of Hypotheses for the 100-kyr Cycle...................................................................29 Martin Claussen, Andre´ Berger and Hermann Held 4. Modelling the 100-kyr Cycle – An Example From LLN EMICs ......................................37 Andre´ Berger and Marie-France Loutre Section 2 Methods of Palaeoclimate Reconstruction and Dating (ed. Frank Sirocko) 5. Introduction – Palaeoclimate Reconstructions and Dating...............................................47 Frank Sirocko 6. Late Quaternary Interglacials in East Antarctica From Ice-Core Dust Records ............53 Barbara Delmonte, Jean Robert Petit, Isabelle Basile-Doelsch, Emil Jagoutz and Valter Maggi 7. Eustatic Sea Level During Past Interglacials .......................................................................75 M. Siddall, J. Chappell and E.-K. Potter 8. Uranium-Series Dating of Peat from Central and Northern Europe...............................93 Manfred Frechen, Melanie Sierralta, Deniz Oezen and Brigitte Urban 9. U-Redistribution in Fossil Reef Corals from Barbados, West Indies, and Sea-Level Reconstruction for MIS 6.5 .................................................................................119 Denis Scholz, Augusto Mangini and Dieter Meischner 10. Holocene and Eemian Varve Types of Eifel Maar Lake Sediments...............................141 Bert Rein, Knut Ja¨ger, Yvonne Kocot, Kirsten Grimm and Frank Sirocko 11. Dating of Interglacial Sediments by Luminescence Methods.........................................157 D. Degering and M.R. Krbetschek 12. Neanderthal Presence and Behaviour in Central and Northwestern Europe During MIS 5e ..........................................................................................................173 Stefan Wenzel

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Section 3 Climate and Vegetation in Europe During MIS 5 (ed. Maria Fernanda Sańchez Gon˜i) 13. Introduction to Climate and Vegetation in Europe During MIS 5.................................197 Marıá Fernanda Sańchez Gon˜i 14. Abrupt Cooling Events at the Very End of the Last Interglacial....................................207 Klemens Seelos and Frank Sirocko 15. Estimates of Temperature and Precipitation Variations During the Eemian Interglacial: New Data From the Grande Pile Record (GP XXI) ...............231 Denis-Didier Rousseau, Christine Hatte´, Danielle Duzer, Patrick Schevin, George Kukla and Joel Guiot 16. Quantitative Time-Series Reconstructions of Holsteinian and Eemian Temperatures Using Botanical Data .............................................................239 Norbert Ku¨hl and Thomas Litt 17. Comparative Analysis of Vegetation and Climate Changes During the Eemian Interglacial in Central and Eastern Europe ..................................................255 A.A. Velichko, E.Y. Novenko, E.M. Zelikson, T. Boettger and F.W. Junge 18. Indications of Short-Term Climate Warming at the Very End of the Eemian in Terrestrial Records of Central and Eastern Europe.......................................265 T. Boettger, F.W. Junge, S. Knetsch, E.Y. Novenko, O.K. Borisova, K.V. Kremenetski and A.A. Velichko 19. Vegetation Dynamics in Southern Germany During Marine Isotope Stage 5 ( 130 to 70 kyr Ago) .................................................................................277 Ulrich C. Mu¨ller and Maria F. Sańchez Gon˜i 20. Subtropical NW Atlantic Surface Water Variability During the Last Interglacial .......289 M.J. Vautravers, G. Bianchi and N.J. Shackleton 21. Abrupt Change of El Ninõ Activity off Peru During Stage MIS 5e-d...........................305 Rein, Bert, Frank Sirocko, Andreas Lu¨ckge, Lutz Reinhardt, Anja Wolf and Wolf-Christian Dullo 22. Interglacial and Glacial Fingerprints from Lake Deposits in the Gobi Desert, NW China ........................................................................................................323 Bernd Wu¨nnemann, Kai Hartmann, Norbert Altmann, Ulrich Hambach, Hans-Joachim Pachur and Hucai Zhang Section 4 Climate, Vegetation and Mammalian Faunas in Europe during Middle Pleistocene Interglacials (MIS 7, 9, 11) (ed. Thomas Litt) 23. Introduction: Climate, Vegetation and Mammalian Faunas in Europe during Middle Pleistocene Interglacials (MIS 7, 9, 11)................................................................................351 Thomas Litt 24. Fine-Tuning the Land–ocean Correlation for the Late Middle Pleistocene of Southern Europe...........................................................................................359 K.H. Roucoux, P.C. Tzedakis, L. de Abreu and N.J. Shackleton

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25. Climate Variability of the Last Five Isotopic Interglacials: Direct Land–Sea–Ice Correlation from the Multiproxy Analysis of North-Western Iberian Margin Deep-Sea Cores ........................................................................................................375 S. Desprat, M.F. Sańchez Gon˜i, F. Naughton, J.-L. Turon, J. Duprat, B. Malaize´, E. Cortijo and J.-P. Peypouquet 26. Palynological and Geochronological Study of the Holsteinian/Hoxnian/Landos Interglacial...............................................................................................................................387 Mebus A. Geyh and Helmut Mu¨ller 27. A New Holsteinian Pollen Record From the Dry Maar at Do¨ttingen (Eifel) ...............397 Markus Diehl and Frank Sirocko 28. Interglacial Pollen Records from Scho¨ningen, North Germany.....................................417 Brigitte Urban 29. Mammalian Faunas From the Interglacial Periods in Central Europe and Their Stratigraphic Correlation....................................................................................445 Wighart von Koenigswald 30. MIS 5 to MIS 8 – Numerically Dated Palaeontological Cave Sites of Central Europe .........................................................................................................455 Wilfried Rosendahl, Doris Do¨ppes and Stephan Kempe 31. The Last and the Penultimate Interglacial as Recorded by Speleothems From a Climatically Sensitive High-Elevation Cave Site in the Alps .........................................471 Christoph Spo¨tl, Steffen Holzka¨mper and Augusto Mangini Section 5 Modelling Past Interglacial Climates (ed. Martin Claussen) 32. Climate System Models – A Brief Introduction................................................................495 Martin Claussen 33. Simulations of the Eemian Interglacial and the Subsequent Glacial Inception with a Coupled Ocean–Atmosphere General Circulation Model..................................499 Frank Kaspar and Ulrich Cubasch 34. Simulated Teleconnections During the Eemian, the Last Glacial Inception and the Preindustrial Period ................................................................................................517 Martin Widmann, Nikolaus Groll and Julie M. Jones 35. Orbital Forcing on Atmospheric Dynamics During the Last Interglacial and Glacial Inception ............................................................................................................527 Gerrit Lohmann and Stephan J. Lorenz 36. Interglacials as Simulated by the LLN 2-D NH and MoBidiC Climate Models..........547 M.F. Loutre, A. Berger, M. Crucifix, S. Desprat and M.F. Sańchez Gon˜i 37. Vegetation–Climate Feedbacks in Transient Simulations Over the Last Interglacial (128 000–113 000 yr BP) .....................................................................................563 M. Gro¨ger, E. Maier-Reimer, U. Mikolajewicz, G. Schurgers, M. Vizcaino and A. Winguth

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38. Mechanisms Leading to the Last Glacial Inception over North America: Results From the CLIMBER-GREMLIN Atmosphere–Ocean–Vegetation Northern Hemisphere Ice-Sheet Model .............................................................................573 Masa Kageyama, Sylvie Charbit, Catherine Ritz, Myriam Khodri and Gilles Ramstein 39. Modelling the End of an Interglacial (MIS 1, 5, 7, 9, 11)..................................................583 Claudia Kubatzki, Martin Claussen, Reinhard Calov and Andrey Ganopolski Section 6 Synthesis 40. Chronology and Climate Forcing of the Last Four Interglacials....................................597 Frank Sirocko, Martin Claussen, Thomas Litt, Maria Fernanda Sańchez Gon˜i, Andre´ Berger, Tatjana Boettger, Markus Diehl, Ste´phanie Desprat, Barbara Delmonte, Detlev Degering, Manfred Frechen, Mebus A. Geyh, Matthias Groeger, Masa Kageyama, Frank Kaspar, Norbert Ku¨hl, Claudia Kubatzki, Gerrit Lohmann, Marie-France Loutre, Ulrich Mu¨ller, Bert Rein, Wilfried Rosendahl, Katy Roucoux, Denis-Didier Rousseau, Klemens Seelos, Mark Siddall, Denis Scholz, Christoph Spo¨tl, Brigitte Urban, Maryline Vautravers, Andrei Velichko, Stefan Wenzel, Martin Widmann and Bernd Wu¨nnemann Index .............................................................................................................................615

Preface: Climates of Past Interglacials – a PAGES Perspective Thorsten Kiefer and Christoph Kull PAGES International Project Office, Sulgeneckstrasse 38, CH-3007- Berne, Switzerland

The cultural evolution of humans has accelerated considerably during the Holocene interglacial. This explosion of civilisation has probably only been possible under the mild and relatively stable climatic conditions that have prevailed for the last 11 000 years. However, these conditions cannot be taken for granted. This is one of the rather simple but unequivocal and important lessons we have learned from the palaeoclimate record. All other interglacials terminated after a few thousands to a few tens of thousands of years. In fact, interglacial states similar to those of today, with little land ice and largely elevated temperatures at mid-high latitudes, prevailed during no more than 15% of the last half million years. These simple empirics already give clear evidence that interglacials are rather unstable on a 10 000-year timescale. Also on shorter timescales of millennia to decades, late Holocene climate fluctuations such as the Little Ice Age and those associated with the Maunder Minimum proove that interglacial climate is not entirely stable on a regional scale but responds to even subtle changes in radiative forcing. Moreover, the discovery of the 8.2 kyr cooling event made it clear that even the worst-case scenario (socioeconomically speaking) of an abrupt change of climate within years is not just a theoretical possibility but has in fact happened in the prehistoric past. It is therefore clear that in principle it could happen again, once some perturbation exceeds a critical threshold of the climate system. Given that humans are indeed becoming increasingly effective at perturbing the climate–environment system, learning

about its sensitivity, thresholds and feedbacks should be in our best interest. An obvious way to do this is to study climate and environment in a suite of experiments where boundary conditions are similar but not identical to those of today. Through the quasi-cyclic reoccurrence of interglacials during the late Pleistocene, we have several such experiments at hand. The palaeoclimate community therefore holds an important key to scientific information on climate change that provides a basis for appropriate adaptation and mitigation strategies. The authors of this book have taken up this challenge and summarise their results in this special volume on climates of past interglacials. It presents state-ofthe-art science on new reconstructions from all spheres of the Earth System and on their synthesis, on methodological advances and on the current ability of numerical models to simulate low- and high-frequency changes of climate, environment and chemical cycling related to interglacials. Most of the authors had been involved in the German DEKLIM programme (www.deklim.de) and have attended some of the five DEKLIM workshops and conferences between 2001 and 2005. The discussions were quite controversial in the beginning, but step by step the dating, reconstruction and modelling of interglacial climate evolution came to a convincing synthesis. Not all open questions are settled, and there is not full consensus on all subjects, but the picture of past climates is now much clearer than it was five years ago. Beyond the pure scientific findings, DEKLIM will leave more (secondary) footprints

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Preface

in palaeoscience which are of equally high value in the perspective of PAGES because of their integrative character. Two fields of palaeoscience that did not often mix well with classic palaeoclimatology were deliberately and successfully incorporated as genuine counterparts of the project. Firstly, studies on palaeovegetation added important information on the regional expression of, and regional feedbacks to, global change. Secondly, numerical modelling of climatic and environmental and geochemical parameters was an integral part of the project. It seems that a generation of modellers and data people has evolved who are capable of communicating effectively on mutual needs and complementary results. A great number of papers of this volume reflect this successful data-model symbiosis. In addition, in spite of being funded by the German Federal Ministry of Education and Research, DEKLIM managed to cooperate successfully beyond national boundaries. The science community and the science consequently benefited greatly from the involvement of many renowned international scientists. The lasting value of this project comes not only from its scientific results, which have led to an improved understanding of the operation of the climate system, but also from the identification of current scientific limitations and open questions. These provide orientation as to where researchers, funders and programmes like PAGES might most effectively focus their resources and efforts for future research. For example, one fundamental limitation is given by the uncertainties in the age control on pre-Holocene interglacials. They are out of reach of radiocarbon dating, not anywhere near the next palaeomagnetic reversal, and the relative smoothness of the climatic records leaves only minor characteristics for event stratigraphy. However, to enable modellers to prescribe forcings with a realistic phasing will require a submillennial chronological accuracy. Creative new approaches are needed, some of which are discussed in this volume.

A similarly high standard is demanded from the quality of the reconstruction of climatic and environmental parameters. The proxies used to reconstruct past interglacial conditions need to detect variations that are usually rather subtle compared to glacial or semi-glacial scenarios. This task imposes higher requirements on methods in terms of precision and accuracy. Hence, the ongoing improvement of analytics, and the development and refinement of proxies and their calibrations, is not just an academic obsession but a fundamental need to advance our understanding of past climate and ultimately the accuracy of prediction. One limitation on model simulations is records of climate forcing, both natural (orbital, solar irradiance, volcanic and greenhouse gases) and anthropogenic (greenhouse gases, aerosols and land use). The validity of any climate simulation depends on the accuracy of the input parameters that generate climate change and environmental response. The abovementioned improvements of chronological tools and proxies will contribute to the advancement in the accuracy of forcing records. Raising the awareness of its importance may be another measure to foster progress. In previous years and decades, a lot of effort was concentrated on understanding the large-amplitude climate variations during glacials, and between glacials and interglacials, thereby disregarding interglacial global change issues. As a result, the spatial density and temporal resolution of records are presently insufficient to allow for firm conclusions on high-frequency or regionalscale climate variability during past interglacials. This becomes more and more acute the further back in time one goes due to the increased difficulty in obtaining good archives. The ambitious target for the coming years will be to reconstruct and simulate centennial-to-decadal scale variability of climate models, not only for the Holocene, but also for the slightly differing boundary conditions of pre-Holocene interglacials.

Preface

Since climate models reveal themselves to be relatively patchy dynamic climate patterns, a critical density of data is necessary to map them. This will require concerted efforts focussing on particular regions and time intervals. The largest region identified as being underrepresented in the climate record and therefore a prime target of future research is the entire southern hemisphere. To be effective, concerted efforts, such as those described above, will require an infrastructure for data compilation, storage and accessibility. This will be important in order to exploit the full potential of numerical climate models. Another fundamental future challenge will be to further increase

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professionalism in both data management and the attitude towards data dissemination. The challenges described above are among those also adopted by PAGES as being particularly worthy of support. These objectives will not only require excellent research by individual scientists and groups but also a high level of interdisciplinarity and organisational structure. While PAGES cannot influence individual scientific excellence, it will continue to facilitate integrative palaeoscience along the lines of the DEKLIM project. Meanwhile, the community can build on many scientific results and datasets, an international network and a comprehensive concept for an integrative global change research programme.


Acknowledgements Our sincere thanks go to Ulrich Katenkamp (Federal Ministry of Education and Research, Germany) for planning and organizing the DEKLIM program, which generously funded research, workshops and a large conference on the climate of past interglacials.

This book could not have been put together without the continuous, meticulous care of Saskia Rudert, who took responsibility for the DEKLIM-EEM project office and for this book in particular.


Section 1 Forcing Mechanisms (ed. Martin Claussen)


1. Introduction to Climate Forcing and Climate Feedbacks Martin Claussen Meteorological Institute, University Hamburg, and Max Planck Institute for Meteorology, Bundesstr. 53, D-20146 Hamburg, Germany

1.1 WHAT IS CLIMATE? When interpreting palaeoclimate archives, the question arises as to what is the driver, or forcing, of climate. For example, when analysing reconstructions of temperature and atmospheric CO2 concentrations during the past interglacials and during the transition from an interglacial to a glacial, it is often asked whether temperature drives changes in atmospheric CO2 concentrations or vice versa – of course with a predisposed assumption of relevance for present-day climate change. Hence the perception of climate forcing, of what is external or internal, is of outmost importance in climate research. To provide some guidance on the discussion in this book, this introductory paper provides a definition of climate, with a discussion on climate forcing and climate feedback and its relevance to the interpretation of palaeoclimate archives. In a classical definition, climate is viewed as the sum of all meteorological phenomena that characterize the mean state of the atmosphere (Hann, 1883), or in more colloquial terms, as mean weather. With the notion of climate variations, not only the mean state of the atmosphere, but also its variability, including the statistics of extreme events, has been incorporated into the definition of climate. Hence in meteorological terms – meteorological, because this definition is made in terms of meteorological state variables – climate is briefly known as the statistics of weather (Hantel et al., 1987 ). The World Meteorological Organization has defined a time span of 30 years over which the statistics of weather is to be calculated, and the climate in the 30-year period of 1931 to 1960 is sometimes referred to as ‘normal climate’. But shorter and longer

time spans can be found in the literature as well – there is nothing special about choosing 30 years. It is required that the time span to define climate has to be choesn such that the statistical moments are stable. However, when browsing through climate archives, it becomes immediately evident that climate is not a stationary process. Climate changes on all timescales. The classical definition of climate has proven to be useful for climatology, the descriptive view of climate. However, for understanding climate dynamics, that is, the processes which govern the mean state of the atmosphere, the classical definition appears to be too restrictive since the mean state of the atmosphere is affected by more than just atmospheric phenomena. It has been realized that the mean state of the atmosphere as well as its variations depends on the dynamics of the climate system and on the interaction between the various components of this system (see Fig. 1.1). von Humboldt (1845) was probably the first who explicitly stated that problem. Therefore in climate dynamics and climate physics, a wider definition of climate has been proposed in terms of state and statistics of the climate system (e.g. Kraus, 2000; Annex in: Houghton et al., 2001). The definition of a climate system is not deduced from first principles. It is done ad hoc, in a pragmatic way to isolate the focus of research, the system, from its external environment. One could define a hierarchy of systems depending on timescales considered (see Peixoto and Oort, 1992). For investigation of seasonal climate variations, for example, the dynamics of the atmosphere, the upper ocean, the terrestrial biosphere

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Martin Claussen

The climate system

Energy

ATMOSPHERE TERRESTRIAL VEGETATION

Water Momentum

Carbon

CRYOSPHERE

OCEAN

MARINE BIOTA

PEDOSPHERE, LITHOSPHERE

© CLIMBER

Fig. 1.1 Sketch of the climate system and its components: the atmosphere, the hydrosphere (mainly the ocean, but also rivers, lakes, rain and groundwater are included), the cryosphere (i.e. ice sheets, glacier, sea ice, snow, permafrost and clathrates), the marine and terrestrial biosphere, the pedosphere (soils and rocks) and – not explicitly shown here – the lithosphere and the upper Earth’s mantle. The latter have to be considered when long times over which the bedrock of ice sheets and the continent changes are considered. The climate system is driven by the solar energy flux and the geothermal heat flux. If Pleistocene climate variations are discussed, volcanic activity and outgasing at oceanic ridges are considered as external forcing. The climate system components are coupled through fluxes of energy, momentum and matter.

(excluding migration of vegetation), sea ice and the upper metres of soil need to be considered. The location of continents, the state of the deep ocean and of the ice sheets can be assumed to be constant in time. Hence, processes in the atmosphere, upper ocean, terrestrial biosphere, sea ice and soil as well as exchange processes, or feedbacks, between these components are internal processes. The heat and mass fluxes from the deep ocean into the upper ocean and from the ice sheets to the atmosphere as well as the geothermal heat flux appear as external forcing. Although a reduction of the climate system to fewer components is often done in climate system modelling, it is problematic; the dynamics of the reduced system can be quite different from the dynamics of the more complete system. In Section 5 of this book (see Chapter 39 in particular) in which evidence from palaeoclimate archives is interpreted by climate system models, these problems will be revisited.

1.2 WHY DOES CLIMATE VARY? If one would sketch climate variability, for example in terms of temperature variance, as a function of frequency, one would obtain a spectrum like the one shown in Fig. 1.2. This figure reveals a so-called red spectrum, that is, an increase in variability with decreasing frequency. Moreover, the red spectrum eventually levels off and merges into a flat, or white, spectrum at low frequencies. Superimposed on the spectrum are a few spikes some of which appear as a direct response to forcing as, for example, the diurnal or the annual cycle. Such a spectrum indicates that climate variability emerges as a direct response to changes in forcing and to some internal processes seemingly independent of a direct forcing. This behaviour has been described as forced or external variability and free or internal variability, respectively (Lorenz, 1979).

Surface temperature variance

Introduction to Climate Forcing and Climate Feedbacks

ORBITAL CYCLES Annual cycle and harmonics

Diurnal cycle and harmonics

Cryosphere Deep-ocean circulation Mixed layer ocean Atmosphere

106

105

104

Thermal relaxation Inertial relaxation 103

102

101

100

10–1

10–2

10–3

10–4

10–5

Period (years)

Fig. 1.2 Sketch of a temperature variance in the nearsurface atmosphere as function of frequency. This figure is taken from Crowley and North (1991) with the permission of Oxford University Press.

1.3 CLIMATE FORCING The most important external forcing is the solar energy flux. This energy flux increases slowly by about 10% per billion years as the Sun becomes steadily hotter. But there are also variations of the solar energy flux at periods of 11, 22, 80 and roughly 200 years. Presumably, these variations amount to only 0.1 to 0.3% of the mean flux, which, however, is only known for the shortest 11-year cycle from satellite measurements. For longer-term variations, estimates are derived from proxy data in combination with theories of solar dynamics (e.g. Lean et al., 1995; Bard et al., 2000). Albeit small, a direct response of the near-surface atmosphere to estimates in solar forcing is found in many climate models, in particular in simulations of the last 1000 years before the industrial revolution (Jones and Mann, 2004). It has also been hypothesized that the atmosphere could react in a nonlinear way to insolation changes by amplifying the weak forcing via wave interaction between the troposphere, the lower 10-km deep ‘weather’ layer of the atmosphere, and the upper atmosphere where the much larger variations in UV radiation modify atmospheric chemistry and hence, the energy budget (e.g. Shindell et al., 1999).

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Solar wind and the magnetic field of the Earth shield the Earth from cosmic radiation. There are a number of studies that reveal correlations between changes in cosmic rays and atmospheric variables (e.g. Svensmark and Friis-Christensen, 1997). In most cases, these correlations appear as artefacts because the correlations vanish if new data appear, or data are simply not treated properly (Laut, 2003). There are several hypotheses of how cosmic rays could affect climate; none of them, however, is commonly accepted (Ramaswamy et al., 2001). The Earth rotates and it spins like a top around the Sun. Since the Sun, the Earth’s moon and the larger planets exert torques on the spinning Earth, the Earth wobbles – the eccentricity of the Earth’s orbit, the obliquity of the Earth’s axis and the perihelion (the position of the Earth closest to the Sun) change. These orbital changes affect the meridional and seasonal distribution of insolation (see Chapter 2). This way, the orbital signal is seen in many climate archives (see also Chapters 3 and 4). Volcanic explosions affect the chemistry of the atmosphere such that a global-scale cooling of some tenths of a degree over two to three years after the explosion can be detected. For example, the Pinatubo eruption cooled the near-surface Earth by some 0.3 K. Hence, it is not the individual volcanic eruption that influences long-term climate variation, but the frequency of eruptions. For example, the lack of eruptions at the beginning of the twentieth century contributed to the observed global warming at that time (Briffa et al., 1998). Volcanic forcing belongs to the class of tectonic forcing. Tectonic forcing is driven by mantle convection, that is, by processes in the Earth’s interior. Whether these processes are defined as forcing or feedback depends on the timescales under consideration. In climate system models that aim to describe the evolution of the Earth’s climate over billions of years, tectonic processes such as spreading of oceanic crust and subduction of continental plates have to be

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considered as internal processes (e.g. von Bloh et al., 2003). For simulations of Pleistocene climate variations, tectonic processes can be considered as external. In some books, humankind is considered as part of the biosphere and, thus, included in the climate system (Kraus, 2000). This approach is, however, not commonly accepted. Simulation of the natural components of the climate system relies on the physics of motion and chemistry of gases and liquids. Human activity, economics, culture and values are, however, not accessible by these tools. Therefore, anthropogenic activity is taken as external forcing in climate system models. The so-called integrated models of the Earth system (including human dynamics as well as the climate system, e.g. Alcamo, 1994; Schellnhuber, 1999) require new approaches. Humankind affects the climate system mainly by altering the chemical composition of the atmosphere and by changing the land-surface structure. Currently, anthropogenic emissions of CO2 by burning fossil fuel and harvesting forests is 50 to 100 times stronger than CO2 emission by outgasing from the Earth’s interior, and the atmospheric CO2 concentration reaches levels of 380 ppmv (parts of CO2 volume per million parts of atmospheric volume) which is well above the preindustrial value of approximately 280 ppmv and the late Pleistocene average of approximately 210 ppmv (Prentice et al., 2001); the atmospheric CH4 concentration rose from approximately 700 ppbv (parts per billion) in the preindustrial to approximately 1750 ppbv, with a late Pleistocene average of approximately 450 ppbv (Prather et al. 2001). Furthermore, a number of new chemical substances with a much stronger potential as a greenhouse gas than CO2 and CH4 have been created by humankind. One-third to one-half of the land is directly affected by land use (Vitousek et al., 1997). Besides CO2 emissions, this leads to changes in the near-surface energy budget, because a deforested area – in particular, when snow

covered – reflects more sunlight than forests and affects transpiration. The brightening of the Earth’s surface due to deforestation has presumably contributed to the cooling during the last centuries by approximately 0.35 K over the last 1000 years (Bauer et al., 2003). 1.4 INTERNAL VARIABILITY Even if there were no changes in external forcing, climate would vary anyhow because of internal instabilities embedded in the climate system components. An illustrative example of this free, internal variability is the motion of a thin layer of oil in a fry pan which is heated from below. If some ingredients are added to visualize the oil flow (pepper, for example), then one can observe regular structures like hexagonal cells or parallel streaks although the heating from below is (ideally) perfectly homogeneous. The structure of flow patterns depends on the strength of the (constant) forcing, that is, on the strength of the imposed heating. With a strong heating, the pattern could exhibit a wavy motion albeit the heating itself is kept steady. Finally, there is a possibility of a turbulent motion of boiling oil – again with a heating which is constant in space and time. Similarly, even with a steady external forcing, the atmosphere would heat at the equator and cool at the poles, and, if the Earth does not rotate, there would be a strong temperature gradient between the day and the night side of the Earth. An atmospheric motion could exist in response to this temperature gradient. Superimposed on a mean circulation between the hot side and the cold side of a resting Earth, atmospheric variability would exist as it does in the case of a homogeneously and constantly heated oil in a fry pan. The weather pattern, of course, would look rather different from our current patterns of zonally moving lowand high-pressure systems. Because the climate system components have different response times or, in other


words, are more of less sluggish, the interaction between them leads to a red spectrum (see Hasselmann, 1976, and Fig. 2 from Crowley and North, 1991). Random variations in a fast climate system component, say the atmosphere, could by chance add up to an anomaly which affects a more sluggish component, the ocean for example. The oceanic random motion generated in this way may also reveal anomalies which then are felt by an even more sluggish ice sheets. Hence, anomalies add up randomly with larger amplitude in the slower variations. Wunsch (2003) argues that this process is the main source of climate variations, while external forcing would add only a little bit (see also Chapter 3).

7

phenomenon of the ‘green’ Sahara. According to palaeoclimatic evidence (e.g. Prentice et al., 2000), the Sahara was much greener in the early and mid-Holocene approximately 9000 to 6000 years ago. The greening has been explained by an increase of the palaeomonsoon in response to changes in the Earth’s orbit which led to a stronger warming of continents (Kutzbach and Guetter, 1986). The resulting increase in monsoonal precipitation appeared to be insufficient to cause any Saharan greening in all climate models. Subsequently, it was shown that a biogeophysical feedback, originally proposed by Charney (1975), could amplify the precipitation to generate enough rain for a substantial northward shift of savannah and steppes (Claussen, 1997).

1.5 CLIMATE AMPLIFIER 1.6 CLIMATE CHANGE TRIGGERS Internal processes and feedbacks between climate system components may not be caused by internal climate variability only. Feedbacks can also amplify external forcing in a disproportional, or nonlinear, way. For example, emission of greenhouse gases in present-day climate would increase the global and annual mean temperature by approximately 1 K if the concentration of greenhouse gases is doubled. This warming leads to larger evaporation and hence an increase in the most important greenhouse gas, water vapour. The so-called moist greenhouse effect amplifies the initial warming. Other feedbacks can further amplify or attenuate the warming. The final warming lies in the range of 1.5–4.5 K. This uncertainty of warming to a doubling in greenhouse gas concentration reflects the uncertainty in understanding feedback processes. Interestingly, the range of uncertainty has not changed very much over the last two decades (McAvaney et al., 2001), but there are attempts underway to reduce this uncertainty by analysing information from palaeoclimate archives. A further illustrative example of positive, that is, amplifying, feedbacks is the

In climate archives, many examples of rapid, abrupt climate changes are found. These could be interpreted as random fluctuations or random jumps of the climate system from one state to another which is one type of internal variability. In many cases, however, it seems plausible to assume that the abrupt climate change is triggered by some external driver. Again the Sahara, in particular the abrupt expansion of the Sahara approximately 5500 years ago (de Menocal, 2000), can serve as an illustrative example. In a coupled atmosphere–vegetation model, Claussen (1997) found that in present-day climate, Northern Africa can exhibit two states: an arid state like today and a ‘green’ state like in the early Holocene. For early Holocene conditions, only the green state seems to exist (Claussen and Gayler, 1997). As the external forcing, in this case the summer insolation over the Northern Hemisphere, declines during the Holocene, the atmosphere–vegetation system over Northern Africa moves from a equilibrium state in the early and middle Holocene to a state with two equilibria today. In theory, it would be possible that

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Northern Africa could still be green, because the green state is an equilibrium, or stable, state. However, dynamically speaking, the green state today is much less stable than the arid state. Hence, any perturbation by some ubiquitous changes in the temperature of the tropical North Atlantic, for example, lets the atmosphere-vegetation system jump from the green into the arid state (Brovkin et al., 1998). Hence, the rapid aridification is not directly driven by orbital forcing; rather, it is triggered. Likewise, the rapid Dansgaard/Oeschger fluctuations are interpreted as transitions between two states of the inter-hemispheric Atlantic Ocean circulation which is driven by large-scale gradients of heating and freshwater fluxes (Ganopolski and Rahmstorf, 2001). Interestingly, the Dansgaard/Oeschger fluctuations occur with a remarkably regular pace. Therefore, it has been suggested that some external driver, which itself is too weak to generate any noticeable change in the climate system, would synchronize these fluctuations (Rahmstorf, 2003). Again, this – yet unknown – pacemaker cannot be considered as forcing, but rather as trigger. 1.7 RELEVANCE TO THE INTERPRETATION OF PALAEOCLIMATE ARCHIVES Climate can vary in response to changes in climate forcing, but it also can vary because of internal instabilities independently of any change in forcing. The part of climate variability which is forced by some external process is easily predictable once the forcing is known. On the contrary, internal climate variability is much less assessible to prediction. Only some statistical properties of internal climate variability can be derived. Hence, the knowledge of what is external and of what is internal becomes important when interpreting palaeoclimate archives in view of climate predictability. In geology and geography (Sirocko, personal communication), climate variations

on timescales much longer than the classical averaging period of 30 years to define climate are sometimes considered as determined by some forcing. This is a fallacy because also the sluggish components of the climate system, such as ice sheets and the deep ocean, exhibit internal, and hence less predictable, variability at longer timescales. Furthermore, internal climate variability could, by chance, emerge as seemingly regular, deterministic oscillation; nonetheless it would be futile, in this case, to search for any oscillating forcing behind it. A problem related to the question of external and internal climate variability is the ‘chicken and egg problem’ of temperature and atmospheric CO2 concentration. It has been suggested that the problem of whether temperature forces atmospheric CO2 concentration or vice versa, whether atmospheric CO2 concentration affects temperature, can be solved by determining the leads and lags between changes in temperature and in greenhouse gases found in palaeoclimate archives. From the dynamical point of view, this approach is futile. Firstly, in oscillating coupled systems the chicken and egg problem cannot be solved by analysing leads and lags. There are numerous examples in which an apparently lagging system drives a seemingly leading system. Secondly and more importantly, temperature and greenhouse gases co-evolve in the climate system by affecting each other. An increase in greenhouse gases leads to a rise in atmospheric temperature, and, in turn, an increase in temperature affects vegetation and upper ocean and thereby the carbon fluxes between the climate system components. Hence, one has to understand temperature and atmospheric CO2 fluctuations as feedback processes within the climate system. Only outgasing of CO2 by tectonic processes and anthropogenic greenhouse gas emissions are considered as climate forcing from this point of view. In Section 1.1.3, several climate forcings such as changes in solar energy flux, in


volcanic activity and changes in the Earth’s orbit around the Sun are discussed. Because there are not yet any robust estimates of solar activity and of volcanic activity during the last interglacials, this forcing has not been considered in numerical simulations presented in this book. Hence, orbital forcing is the main agent in the current discussion of the dynamics of past interglacials and glacial inception, while the role of solar and volcanic activity has still to be explored. In general, climate variations appear as both externally driven and internally generated. In particular, internal feedbacks can amplify external forcing which itself would be too weak to cause any detectable climate change. In such cases, forcing should be viewed as trigger, and the relation between trigger and climate change is likely to be highly nonlinear. For example, the idea of a climate change trigger is used to explain glacial inceptions (Calov et al., 2004; see also Chapter 39). In the numerical model of the climate system used by Calov et al. (2004), a glacial inception appears as an instability of the atmosphere–ice sheet system. Once this instability is triggered, then the resulting climate changes are amplified by fast climate feedbacks such as changes in atmospheric water vapour and eventually by slower feedbacks such as vegetation shift, shift in oceanic circulation and in the carbon cycle which, in turn, affects atmospheric CO2 concentration and other greenhouse gases. The analysis of external and internal climate processes is complicated by the fact that, partly due to the lack of computational resources and partly due to the gaps in understanding the climate system, many climate system models do not simulate the interaction between all climate system components explicitly. Instead, only the dynamics of some subsystems, such as atmosphere and ocean for example, is explicitly considered, while the state of the other subsystems, such as the big ice sheets, is prescribed from palaeoclimate archives or other theoretical estimates. As mentioned in Section 1.1.1, the reduction of models of the

9

climate system has to be done with caution, and hence, this method will be critically assessed in Section 5.8. Finally, it is understood that the study of past interglacials and glacial inceptions by both excavating palaeoclimate archives and interpreting palaeoclimatic evidence by a spectrum of climate system models (see Chapter 32) – as done in this book – will considerably advance our understanding of the dynamics of the climate system in response to natural as well as anthropogenic forcing.

ACKNOWLEDGEMENT The author wishes to thank Claudia Kubatzki, Alfred-Wegener Institute, Bremerhaven, and Frank Sirocko, University of Mainz, for constructive discussion, and Saskia Rudert, University Mainz, Ursula Werner, Potsdam Institute for Climate Impact Research, and Barbara Zinecker, Max Planck Institute for Meteorology, for technical and editorial assistance.

REFERENCES Alcamo, J. (ed.), 1994. IMAGE 2.0: Integrated modeling of global climate change. Special issue of Water Air Soil Pollution, 76, 1–321. Bard, E., Raisbeck, G., Yiou, F., Jouzel, J., 2000. Solar irradiance during the last 1200 years based on cosmogenic nuclides. Tellus, 52B, 985–992. Bauer, E., Claussen, M., Brovkin, V., Hu¨nerbein, A., 2003. Assessing climate forcings of the Earth system for the past millennium. Geophysical Research Letter, 30, 1276–1279. von Bloh, W., Bounama, C., Franck, S., 2003. Cambrian explosion triggered by geosphere-biosphere feedbacks. Geophysical Research Letter, 30, 1963–1966. Briffa, K.R., Jones, P.D., Schweingruber, F.H., Osborn, T.J., 1998. Influence of volcanic eruptions on Northern Hemisphere summer temperature over the past 600 years. Science, 393, 450–455. Brovkin, V., Claussen, M., Petoukhov, V., Ganopolski, A., 1998. On the stability of the atmospherevegetation system in the Sahara/Sahel region. Journal of Geophysical Research, 103, 31613–31624.

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Calov, R., Ganopolski, A., Petoukhov, V., Claussen, M., Greve, R., 2004. Transient simulation of the last glacial inception. Part I: Glacial inception as a bifurcation in the climate system. Climate Dynamics, 24(6), 545–562. Charney, J.G., 1975. Dynamics of deserts and droughts in the Sahel. Quarterly Journal of the Royal Meteorological Society, 101, 193–202. Claussen, M., 1997. Modelling biogeophysical feedback in the African and Indian Monsoon region. Climate Dynamics, 13, 247–257. Claussen, M., Gayler, V., 1997. The greening of Sahara during the mid-Holocene: results of an interactive atmosphere – biome model. Global Ecology and Biogeography Letters, 6, 369–377. Crowley, T., North, G., 1991. Paleoclimatology. Oxford Monographs on Geology and Geophysics, 18. New York: Oxford University Press, 339 pp. Ganopolski, A., Rahmstorf, S., 2001. Simulation of rapid glacial climate changes in a coupled climate model. Nature, 409, 153–158. Hann, J. 1883. Handbuch der Klimatologie. Engelhorn, Stuttgart. Hantel, M., Kraus, H., Scho¨nwiese, C.-D., 1987. Climate definition. In: Fischer, G. (Hrsg.): Climatology. Landolt-Bo¨rnstein, Functional Relationships in Science and Technology V/4/c1. Berlin: Springer. Hasselmann, K., 1976. Stochastic models. I. Theory. Tellus, 28, 473–485. Houghton, J.T., Ding, Y., Griggs, D.J., Noguer, M., van der Linden, P., Dai, X., Maskell, K., Johnson, C. I. (eds.), 2001. Climate Change 2001: The Scientific Basis. Contribution of Working Group I to the Third Assessment Report of the Intergovernmental Panel on Climate Change. Cambridge University Press, 881 pp. von Humboldt, A., 1845. Kosmos, Entwurf einer physischen Weltbeschreibung. Band I, J.G. Cotta, Stuttgart and Tu¨bingen, 493 pp. Jones, P.D., Mann, M.E., 2004. Climate over past millennia. Reviews of Geophysics, 42, 1–42. Kraus, H., 2000. Die Atmospha¨re der Erde. Eine Einfu¨hrung in die Meteorologie. Vieweg. Braunschweig/Wiesbaden, 470 pp. Kutzbach, J.E., Guetter, P.J., 1986. The influence of changing orbital parameters and surface boundary conditions on climate simulations for the past 18 000 years. Journal of Atmospheric and Oceanic Science, 43, 1726–1759. Laut, P., 2003. Solar activity and terrestrial climate: some dubious correlations. Journal of Atmospheric & Solar-Terrestrical Physics, 65, 801–812. Lean, J., Beer, J., Bradley, R., 1995. Reconstruction of solar irradiance since 1610: implications for climate change. Geophysical Research Letter, 22, 3195–3198.

Lorenz, E.N., 1979. Forced and free variations of weather and climate. Journal of Atmospheric and Oceanic Science, 36, 1367–1376. de Menocal, P.B., Ortiz, J., Guilderson, T., Adkins, J., Sarnthein, M., Baker, L., Yarusinski, M., 2000. Abrupt onset and termination of the African Humid Period: Rapid climate response to gradual insolation forcing. Quaternary Science Review, 19, 347–361. McAvaney, B.J., Covey, C., Joussaume, S., Kattsov, V., Kitoh, A., Ogana, W., Pitman, A.J., Weaver, A.J., Wood, R.A., Zhao, Z.-C., 2001. Model evaluation. In: Houghton, J.T., Ding, Y., Griggs, D.J., Noguer, M., van der Linden, P., Dai, X., Maskell, K., Johnson, C.I. (eds.), 2001: Climate Change 2001: The Scientific Basis. Contribution of Working Group I to the Third Assessment Report of the Intergovernmental Panel on Climate Change . Cambridge University Press, 881 pp. Peixoto, J.P., Oort, A.H., 1992. Physics of Climate. New York: American Institute of Physics, 111– 125. Prather, M., Ehhalt, D., Dentener, F., Derwent, R., Dlugokencky, E., Holland, E., Isaksen, I., Katima, J., Kirschoff, V., Matson, P., Midgley, P., Wang, M., 2001. Atmospheric chemistry and greenhouse gases. In: Houghton, J.T., Ding, Y., Griggs, D.J., Noguer, M., van der Linden, P., Dai, X., Maskell, K., Johnson, C.I. (eds.), Climate Change 2001: The Scientific Basis. Contribution of Working Group I to the Third Assessment Report of the Intergovernmental Panel on Climate Change. Cambridge University Press, 881 pp. Prentice, I.C., Jolly, D., BIOME 6000 members, 2000. Mid-Holocene and glacial-maximum vegetation geography of the northern continents and Africa. Journal of Biogeography, 27, 507–519. Prentice, I.C., Farquhar, G.D., Fasham, M.J.R., Goulden, M.L., Heimann, M., Jaramillo, V.J., Kheshgi, H.S., Le Que´re´, C., Scholes, R.J., Wallace, D.W.R., 2001. The carbon cycle and atmospheric carbon dioxide. In: Houghton, J.T., Ding, Y., Griggs, D.J., Noguer, M., van der Linden, P., Dai, X., Maskell, K., Johnson, C.I. (eds.), Climate Change 2001: The Scientific Basis. Contribution of Working Group I to the Third Assessment Report of the Intergovernmental Panel on Climate Change Cambridge University Press, 881 pp. Rahmstorf, S., 2003. Timing of abrupt climate change: a precise clock. Geophysical Research Letter, 30, 1510–1513. Ramaswamy, Y., Boucher, O., Jaigh, J., Hauglustaine, D., Haywood, J., Myhre, G., Nakajima, T., Shi, G. Y., Solomon, S., 2001. Radiative forcing of climate change. In: Houghton, J.T., Ding, Y., Griggs, D.J., Noguer, M., van der Linden, P.,

Introduction to Climate Forcing and Climate Feedbacks Dai, X., Maskell, K., Johnson, C.I. (eds.), Climate Change 2001: The Scientific Basis. Contribution of Working Group I to the Third Assessment Report of the Intergovernmental Panel on Climate Change. Cambridge University Press, 881 pp. Schellnhuber, H.J., 1999. ‘Earth system’ analysis and the second Copernican revolution. Nature, 402, C19–C26. Shindell, D., Rind, D., Balachandran, N., Lean, J., Lonergan, P., 1999. Solar cycle variability, ozone, and climate. Science, 284, 305–308.

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Svensmark, H., Friis-Christensen, E., 1997. Variation of cosmic ray flux and global could coverage – a missing link in solar–climate relationships. Journal of Atmospheric and Solar-Terrestrial Physics, 59, 1225–1232. Vitousek, P.M., Mooney, H.A., Lubchenco, J., Melillo, J.M., 1997. Human domination of Earth’s ecosystems. Science 277, 494–499. Wunsch, C., 2003. The spectral description of climate change including the 100ky energy. Climate Dynamics, 20, 353–363.


2. Insolation During Interglacial A. Berger1, M.F. Loutre1, F. Kaspar2 and S.J. Lorenz3 1

Universite´ catholique de Louvain, Institut d’Astronomie et de Geóphysique G. Lemaıˆtre, 1348 Louvain-la-Neuve, Belgium 2 Institut fu¨r Meteorologie, Freie Universita¨t Berlin, 12165 Berlin, Germany 3 Meteorologisches Institut, Universita¨t Hamburg, 20146 Hamburg, Germany

ABSTRACT The main insolation parameters are reviewed, in particular the energy received by the whole Earth over a full year and the 24-hour mean irradiance. Their spectral characteristics are underlined, and some remarks are made about the differences between insolation at the top of the atmosphere and at the Earth’s surface, the caloric and the astronomical seasons, the insolations in the tropical and in the high latitudes, the mid-month and the calendar insolations. The insolations characterizing interglacials marine isotope stage (MIS) 1 to MIS 11 are described, and their common features are discussed. 2.1 INTRODUCTION A link between climate and insolation forcing is indicated by numerous records of past environmental changes. Adhe´mar (1842) suggested that variations in the orbital configuration of the Earth, more specifically the precession of the equinoxes, could be responsible for glacial–interglacial cycles. In the second half of the nineteenth century, Croll (1875) approached the glaciation problem from the synergistic standpoint of the combined effects of all three major astronomical factors on seasonal insolation during perihelion and aphelion. A specific characteristic of his model essentially lies in his hypothesis that the critical season for the initiation of glacial stages is northern hemisphere (NH) winter. He argued that a decrease in the amount of

sunlight received during the winter favours the accumulation of snow and that any small initial increase in the size of the area covered by snow would be amplified by the snowfields themselves (positive feedback). Almost at the same time, Murphy (1876) adopted the opposite view that a long, cool summer and a short, mild winter provide the most favourable conditions for glaciation, a theory later recognized by Bru¨ckner et al. (1925). Today, the astronomical theory of palaeoclimates is mainly associated with the name of Milankovitch. Milankovitch (1941) was among the firsts to argue that the insolation during the NH summer could be the critical parameter and to take into account the effects of all astronomical parameters (eccentricity, obliquity and position of the perihelion). He integrated their effects to compute the insolation at the top of the atmosphere. Actually, Milankovitch (1941, p. 542 original edition) cites Ko¨ppen who ‘after an exhaustive discussion of all the possibility, . . . answered the question by indicating that it is the diminution of heat during the summer half-year which is the decisive factor in glaciation’ (p. 544 English edition). This discussion can be found in Ko¨ppen and Wegener (1924) where Ko¨ppen acknowledges the work of Penck and Bru¨ckner (1909) who suggested that glaciation is a problem not of enhanced precipitation, but of reduced ablation. An overview of the development of astronomical theories of palaeoclimate is given by Berger (1988) and by Paillard (2001). Algorithms for the precise calculation of the insolation patterns have been published previously (e.g. Berger,

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1978). Here, we present some fundamental reflections on the underlying mechanisms and discuss differences and similarities between the interglacials of the past 500 000 years. The article is supplemented with digital material containing tables with time series of insolation and plots for the interglacials.

solar constant, S0, the semi-major axis, a, of the Earth’s orbit around the Sun (the ecliptic), of its eccentricity, e, its obliquity, ", and the longitude of the perihelion measured ~ (Fig. 2.1). from the moving equinox, ! S0 is calculated at the mean distance, rm , from the Earth to the Sun. On the energy point of view, rm is given by: pffiffiffiffiffiffiffiffiffiffiffiffi r 2m ¼ a2 1 e2

2.2 ENERGY RECEIVED BY THE WHOLE EARTH OVER A FULL YEAR Assuming a perfectly transparent atmosphere, the energy, W, available at any given latitude on the Earth at any time during the year is a single-valued function of the

Therefore: Sa S0 ¼ pffiffiffiffiffiffiffiffiffiffiffiffi 1 e2

Sun Earth

Variation of the ‘eccentricity’ Today Summer solstice 22 June

Spring equinox 21 March 92.8 89.0

23°27′

Sun

93.6 89.8 Fall equinox 23 September

Winter solstice 22 December

Fig. 2.1 Top: the shape of the Earth’s orbit around the Sun is given by the eccentricity (it must be kept in mind that the semi-major axis is an invariant). Bottom left: The climatic precession parameter and the present-day configuration of the Earth’s orbit, with the winter solstice occurring very close to the perihelion. The length of the seasons is also indicated in days. Bottom right: the tilt of the Earth’s axis of rotation is called obliquity, and its present-day value is 23 279. These figures are reproduced from ‘European Science Foundation European Latsis Prize 2001’ and extended.

Insolation During Interglacial

where Sa is the energy received by unit of time on a unit area perpendicular to the Sun’s rays situated at a distance a from the Sun (Berger and Loutre, 1994). As a has no purely secular variations, Sa is only a function of the solar output. In the calculation of the long-term climatic variations over the Quaternary, Sa is supposed to be a constant equal to 1365 Wm2 in this paper.* e is a measure of the shape of the Earth’s orbit around the Sun. Its long-term variations are characterized by a mean period of about 100 kyr superimposed on a longer period of about 400 kyr, although their spectral structure is more complex. Its present-day value is equal to 0.016 and it varies between a maximum of 0.075 and a minimum of zero (the orbit is in this case circular) over the last millions of years. As a consequence of such a range of variation, S0 varies between 1365 and 1369 Wm2 for respectively the maximum and minimum values of e. From this, it follows that the energy received per unit area of the Earth’s surface and per unit of time over a full year is given by

WE ¼

S0 Sa ¼ pffiffiffiffiffiffiffiffiffiffiffiffi 4 4 1 e2

and varies by 1 Wm2 between the extrema of e. e is actually the only astronomical parameter which can change the total energy received by the whole Earth over one year. The other two parameters are redistributing the energy among latitudes and seasons. 2.3 PRECESSION AND OBLIQUITY Although the direct impact of e is rather small, e plays another much more A constant value of 1365 Wm2 is suggested by the Paleo Model Intercomparison Project (PMIP, phase 2) and is used in various modelling studies. The observed value during the last decades was slightly higher, but latest publications concluded that solar activity during the last decades was exceptionally high compared to the previous centuries (Fro¨hlich and Lean, 2004) and even millennia (Solanki et al., 2004). *

15

important role by modulating the amplitude of the climatic precession parameter, ~ This parameter is a measure of the e sin !. Earth–Sun distance at a fixed given time ~ in the year. The present-day value of !, following the calculation by Berger (1978) ~ is measured from the vernal where ! point, is 102 which means that the winter (December) solstice occurs close to the perihelion. This particular feature makes that the winter (summer) in the northern latitudes receives more (less) energy than usual which pleads in favour of using a long-term average for the insolation values instead of the present-day ones as reference. The average period of the long~ is 21 kyr resultterm variations of e sin ! ing from four main spectral components situated around 23 and 19 kyr. This parameter has clearly an opposite effect in the two hemispheres during their similar local seasons. About 12 kyr ago, when the summer solstice was at the perihelion, all the latitudes over the Earth received daily more energy than now during the astronomical NH summer (a season defined between the spring and fall equinoxes). This means that the NH latitudes received more energy during their summer and less during their winter, the reverse being equally true for the southern hemisphere (SH) latitudes during the SH seasons. The seasonal gradient was therefore magnified in the NH and reduced in the SH. " is the tilt of the Earth’s rotational axis relative to a perpendicular drawn to the plane of the ecliptic. Its present-day value is 23 279. It varies steadily with an average period of 41 kyr. An increase of " leads to an insolation increase in all latitudes during their respective summer and the reverse during winter. The seasonal gradient is therefore larger everywhere over the Earth. A decrease of " leads similarly to a decrease in insolation in the summer hemisphere and an increase in the winter hemisphere. As the strength of this effect is small in the tropics and maximum at the poles, it damps the

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seasonal cycle in the high latitudes of both hemisphere simultaneously. As glaciations occur also more or less in phase in both hemispheres, this is most probably one of the reasons for which the obliquity signal appears so clearly in most palaeoclimatic records in high latitudes. The spectral characteristics of the astronomical elements and of the insolations date back from the 1970s only. Emiliani (1955) was one of the firsts to estimate the mean periods of the astronomical parameters by counting the number of peaks from the Milankovitch curves. This led him with 92, 40 and 21 kyr for respec~ Hays et al. (1976) tively e, " and e sin !. used a spectral analysis technique which they applied on the numerical values of the astronomical parameters re-calculated by Brouwer and van Woerkom (1950) and Vernekar (1972). They found 125 and 96 kyr for e, 41 for " and 23 and 19 kyr for precession. In the mean time, Berger (1978) had completed his calculation, providing for the first time a full list of the periods characterizing the theoretical expansion of e (with periods of 400, 125, 95, 99, 131 and 2035 kyr), of " (with peri~ ods of 41, 54 and 29 kyr) and of e sin ! (with periods of 23.7, 22.4, 18.9 and 19.2 kyr) (see Berger, 1978 for a full list). It is the splitting of the 21-kyr precessional period into 23 and 19 kyr found independently in geological and astronomical data which was one of the first most delicate and impressive tests of the Milankovitch theory.

2.4 24-h MEAN IRRADIANCE ~ on the From these impacts of e, " and e sin ! total and the latitudinal and seasonal distributions of the energy received by the Earth from the Sun, the Milankovitch condition for entering into glaciation (1941) can be expressed in terms of orbital parameters. A minimum summer insolation in high

northern latitudes is indeed reached when (i) " is minimum, the seasonal gradient ~ is maxbeing then reduced in NH, (ii) e sin ! imum which implies that both the summer solstice occurs at aphelion (as it is more or less the case now) and e is maximum making the Earth–Sun distance at summer solstice, rA , even larger. As we have rA ¼ að1 þ eÞ; rP ¼ að1 eÞ; rP rA ¼ 2ae; rP being the Earth–Sun distance at perihelion, the difference in the total energy received by the Earth between perihelion and aphelion is proportional to four times the eccentricity. For a maximum of e (0.075), this means a difference reaching 30%; it is presently 6.4%. The Milankovitch hypothesis is also coherent with a warm winter allowing more precipitations over the continental high latitudes. However, a maximum eccentricity has a counter effect, although small, in increasing the total energy received by the Earth. All these considerations can be mathematically deduced from the theoretical calculation of the energy received from the Sun on a horizontal surface for each day and each latitude, . For the latitudes where there is a sunrise and a sunset every day ðjj < 2 jjÞ, the average energy over a time period of 24 h centred on solar noon (the 24-h mean irradiance) and further called incoming solar radiation (insolation) is given by: Sa a 2 W¼ ðH0 sin sin þ cos cos sin H0 Þ r where , the declination of the Sun, is related to its longitude, , by sin ¼ sin sin " being the angular distance from the spring equinox in the ecliptic;


H0 is its hour angle, H, at sunset (at sunrise H ¼ H0 and at solar noon H ¼ 0). It is given by: cos H0 ¼ tan tan r is the distance from the Earth to the Sun given by r¼

að1 eÞ ; 1 þ e cosð !Þ

the numerical value of ! being equal to ~ þ 180 as discussed in Berger et al. (1993). ! r, and are assumed to be constant over one day. For all latitudes jj 2 jj (a) where there is a long polar night: < 0 W ¼0 Length of day ¼ 0 (b) where there is a long polar day: > 0 2 W ¼ S a sin sin r Length of day ¼ 24 hours Analysis of the spectral characteristics of W shows that for a given , (a/r)2 is a function of precession only and the third factor ðH0 sin sin þ cos cos sin H0 Þ is a function of " only through the declination ðÞ and the length of the day (2 H0). The mathematical expression of W allows to conclude (Berger et al., 1993) that its variations are (i) dominated by variations in precession (except close to the polar night), with the obliquity playing a secondary role but more important in high than in low latitudes; (ii) at the equinoxes, insolation is only a function of precession for all latitudes; (iii) at the solstices, both precession and obliquity

17

influence W , but precession is dominating at all latitudes (except close to the polar night). 2.5 INSOLATION AT THE TOP OF THE ATMOSPHERE AND AT THE EARTH’S SURFACE The values of W are usually referred to as ‘insolation at the top of the atmosphere’ because the perturbations by the atmosphere are not taken into account. Their pattern differs however from the pattern of radiation reaching the Earth’s surface, particularly in high latitudes where the surface albedo is large. Using a radiative, convective model, Tricot and Berger (1988) showed that the atmospheric attenuation essentially reduces the absolute variations of the incident solar radiation at the Earth’s surface as compared with the variations of the insolation at the top of the atmosphere. These variations are maximum in high latitudes, but the variations of the absorbed radiation at the Earth’s surface are maximum in tropical and middle latitudes related to the increase of the surface albedo with latitude. In the summer hemisphere, the largescale gradient of insolation between the tropics and the polar regions shows deviations from its present-day value the characteristic frequencies of which depend upon the type of insolation considered: (i) for the extraterrestrial insolation, the main frequency of the variations of the large-scale latitudinal gradient is about 40 kyr, whereas (ii) for the incident and mainly the absorbed insolation at the surface the large-scale gradient shows, in addition, quasi-periodicity of about 23 kyr; this difference is related to the atmospheric attenuation which reduces more strongly the variations of insolation at the surface in high latitudes than in tropics, preventing the obliquity signal to appear.

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2.6 CALORIC SEASONS This behaviour differs from those of the caloric insolations (i.e. the total solar energy received during a half-year caloric season) introduced by Milankovitch (1941). Such a caloric season is exactly half-a-year long, and the caloric summer counts all days for which insolation is larger than for any day of the caloric winter. These caloric insolations are mainly a function of precession in low latitudes and obliquity in high latitudes. Milankovitch’s idea of introducing the caloric seasons was to attempt at solving the difficulty of having to deal with both the total energy received during a given season and its length, which both vary in time. However, it is far from being that easy. The caloric seasons by Milankovitch raise indeed two problems: first, their start changes with time; second it is impossible to define them in the equatorial regions where the seasonal march of insolation has two minima and two maxima. 2.7 TROPICAL LATITUDES In the intertropical belt, the seasonal cycle of insolation is particular because the Sun comes overhead twice a year at each latitude (Berger and Loutre, 1997). At the equator for example, the maxima and the minima are reached at about the equinoxes and the solstices respectively (this is not exactly true because the Earth–Sun distance modulates the effect of the declination, but we may assume that the approximation is acceptable). Moreover, because of precession, the 24-h mean irradiance at the spring equinox will alternatively be larger and lower than at the fall equinox. The same holds for the solstice minima. If we assume that the climate responds to the equatorial absolute maximum, whether it occurs at spring or fall equinox, a 100-kyr (related to eccentricity) and a 11-kyr (half the precession period) appear very strongly in the long-term

variations of the equatorial insolation (Berger et al., 2004). Moreover, the ‘seasonal gradient’ (difference between the absolute maximum and the absolute minimum) introduces a 5.5-kyr periodicity related to one-fourth of the precessional cycle. 2.8 THE ANNUAL AND SEASONAL IRRADIATION The annual irradiation, that is, the total amount of solar energy received during one year at a given latitude does not depend on climatic precession. By adding all the daily values, one must take into account the fact that the Earth–Sun distance varies over the year (by up to 15% for the largest eccentricity value, see Section 2.3). But on the other hand, according to Kepler’s laws, the Earth travels faster at the perihelion than at the aphelion. Integrating over time, these two effects exactly compensate one another such that the total annual irradiation no longer depends on the Earth–Sun distance, but only on the inclination of the Earth’s axis of rotation with respect to the ecliptic, it means on obliquity. In both the northern and southern high latitudes, the total annual irradiation varies in phase with obliquity. Over the last 500 kyr, at 80 N, the maximum amplitude of the variation is 500 106 J m2 around a mean value of 5600 106 J m2 . At low latitudes, the total annual irradiation and obliquity are exactly out of phase. The phase reversal between high and low latitudes varies in time, occurring actually between 43 and 44 (N and S). At the equator, the maximum amplitude of the signal over the last 500 kyr is only 120 106 J m2 , and the mean value is 13 100 106 J m2 . In both the high and low latitude cases, the obliquity variation explains nearly 99% of the variance, the rest being related to eccentricity. Likewise, the total irradiation (in J m2 ) received at a given latitude between two given orbital positions of the Earth, defined


by their true longitude, depends only on obliquity and eccentricity. Seasonal irradiations (astronomical or meteorological) are particular cases. For example, the total summer (JJA) irradiation at 80 N varies over the last 500 kyr in phase with obliquity with an amplitude of 250 106 J m2 , i.e. 9% of its mean value. At the equator, the insolation signal is in phase opposition with obliquity and its amplitude is only 30 106 J m2 , i.e., 0:9% of its mean value. However, the duration of a time interval through the year, a season in particular, is a function of precession. The astronomical seasons are the periods during which the Sun traverses the quadrants of the ecliptic, counted from the vernal equinox. Since the areas of the corresponding sectors are not equal, and since the radius vector of the Sun sweeps out equal areas in equal time, the four seasons have different durations. As an example, over the last 500 kyr, the length of astronomical summer varies between 83.6 and 99.8 days. Consequently, the average insolation over a season (in W m2 ) depends both on precession and on obliquity. However, " being developed around a constant value, the ratio between obliquity and precession is dominated by precession to the first order of magnitude. This leads to a summer mean irradiance at the equator strongly dominated by precession. The power of the obliquity component increases towards the high latitudes, but precession remains the dominant component of the summer mean irradiance for all latitudes. 2.9 MID-MONTH AND CALENDAR INSOLATIONS The longitude varies from 0 (assumed to define the time of spring equinox, which is arbitrarily fixed at March 21 if a calendar has to be used) to 360 , with ¼ 90 , 180 and 270 defining respectively the summer solstice, fall equinox and winter solstice.

19

¼ 0, 30 , 60 . . . define the so-called midmonth values which are actually for around the 20th day of each month. Because the length of the astronomical seasons varies in time (Berger and Loutre, 1994), these midmonth values are not related to a fixed calendar date. If such calendar dates need to be used, the mean longitude m must be used which is linked to through a formula ~ In order to illustrate involving e and !. the difference between the mid-month and the calendar values, a comparison is first made between the 60 N daily insolation at present and 10 000 years before present (BP), for ¼ 210. The difference between these ‘‘mid-month’’ insolations amounts to 5 Wm2 , but ¼ 210 presently refers to 24 October ð109 Wm2 Þ and, at 10 000 years BP, it referred to 16 October ð104 Wm2 Þ. Thus, the difference reflects mainly the secular changes of both obliquity and shape of the ecliptic. Second, if a calendar date insolation is considered, the long-term variations of the length of the astronomical seasons is explicitly recognized. For the same latitude and years, if daily insolations at 16 October are compared, a difference of 29 Wm2 is found. This is a result of the fact that on 16 October the true longitude of the Earth is presently 202 ð133 Wm2 Þ and, at 10 000 years BP, 210 (104 Wm2 ). Another example, Fig. 2.2c and 2.2d illustrates the insolation anomaly at 125 kyr BP relative to today’s conditions. In Fig. 2.2c, anomalies are calculated for the same position on the orbit, that is, the same true solar longitude. In Fig. 2.2d, the anomalies are calculated based on the classical calendar. For both figures, vernal equinox has been used as a reference. Therefore, differences between both figures are small around March. Differences are largest in autumn. At 125 kyr BP, the autumnal equinox is at 12 September, whereas it is at 23 September today. This difference of 11 days leads to significantly shorter summers (85 days 125 kyr ago against 94 today). The largest differences

20

A. Berger et al. 90N

90N

480

60N

60N 3

480

30N

30N

460 440 EQ

420 400 375 350 300 250 200 150 100 50 5

440

30S

60S

440 –15 EQ

440

–9

30S –3

–3

60S

90S

90S –60 Jan

(a)

–30 Feb

0 30 60 90 120 150 180 (VE) Apr May Jun Jul Aug Sep True longitude on Earth’s orbit from vernal equinox (VE)

210 Oct

240 Nov

(b)

50 100 150 200 250 300 350 375 400 420 440 460 480 500 520 540 560 580 600

5

–60 Jan

270 Dec

90N

–30 Feb


–42–39 –36–33–30–27–24–21–18–15–12–9–6–3 3

6

210 Oct

240 Nov

270 Dec

9 12 15 18 21 24 27 30 33 36 39 42

90N

60

–10

60N

–10

60N

–40

–20 –20

30N

–30

30N –30

50 EQ

EQ

40 30

30S –50

30S

20

60S

90S

–60 Jan

(c)

–30 Feb


210 Oct

240 Nov

270 Dec

90S

60

–60

60S

10

0

30 Jan

(d)

–65 –60 –55 –50 –45 –40 –35 –30 –25 –20 –15 –10 –5 5 10 15 20 25 30 35 40 45 50 55 60 65

60 Feb

90 Mar

120 150 180 210 240 270 300 330 360 Apr May Jun Jul Aug Sep Oct Nov Dec (annual cycle)

–65 –60 –55 –50 –45 –40 –35 –30 –25 –20 –15 –10 –5 5 10 15 20 25 30 35 40 45 50 55 60 65

90N

90N

–40 60N

60N

–10

10 –30

60

30N

30N

50 EQ

–20 EQ

40 30

30S

20

10

10

60S

60S

90S

90S –60 Jan

–30 Feb

0 (VE)

30 Apr

60 May

90 Jun

120 Jul

150 Aug

180 Sep

210 Oct

240 Nov

270 Dec

–65 –60 –55 –50 –45 –40 –35 –30 –25 –20 –15 –10 –5 5 10 15 20 25 30 35 40 45 50 55 60 65

10

–60 Jan

–30 Feb

0 (VE)

30 Apr

60 May

90 Jun

120 Jul

150 Aug

180 Sep

210 Oct

240 Nov

270 Dec

True longitude on Earth’s orbit from vernal equinox (VE)

True longitude on Earth’s orbit from vernal equinox (VE)

(e)

–10

30S

20

(f)

–65 –60 –55 –50 –45 –40 –35 –30 –25 –20 –15 –10 –5 5 10 15 20 25 30 35 40 45 50 55 60 65

Fig. 2.2-a (left) Present-day insolation as absolute values. Fig. 2.2-b (right): Present-day insolation as deviation from the mean of the last 800 000 years. Fig. 2.2-c (left) Insolation at 125 kyr as deviation from the present-day value calculated by comparing same positions on the orbit with a fixed vernal equinox. Fig. 2.2-d (right): Same as (2-c), but calculated by comparing same calendar dates with a fixed vernal equinox. Insolation anomaly at 128 kyr Fig. 2.2-e (left) and at 115 kyr Fig. 2.2-f (right) as deviation from the present-day values. Fig. 2.2-a to Fig. 2.2-f Distribution of insolation [Wm2] over latitudes and date in the year. Except for Fig. 2.2-d, anomalies are calculated based on a comparison of the same position on the Earth’s orbit (measured from vernal equinox). Anomalies in Fig. 2.2-c to Fig. 2.2-f refer to today’s values.


2.10 CALENDAR DEFINITION AND PALAEOCLIMATE SIMULATIONS The decision how monthly or seasonal differences are calculated is of particular importance when results of climate simulations are analysed as anomalies from today’s climate. Joussaume and Braconnot (1997) discussed this problem for simulations of the climate at 126 000 yr BP with an atmosphere general circulation model (GCM). They compared simulated monthly temperatures over Europe and precipitation over central Africa using the calendar and mid-month definitions with a vernal equinox fixed at 21 March. The annual cycle showed similar features for both definitions, but with important differences in autumn. These differences between the two definitions were as large as the differences between 126 kyr BP and the present. As autumn started earlier at 126 kyr BP, it was already colder for the same calendar date over Europe and precipitation was weaker over Africa. Large September temperature anomalies due to the definition of the calendar were visible over large parts of NH’s continents, reaching a maximum of more than 10 C over eastern Asia. This difference was larger than the internal variability of the model and exceeds the differences between 126 kyr BP and today significantly. The differences observed by Joussaume and Braconnot (1997) are not fully caused

by direct consequences of the calendar definition, but are partly attributable to a calendar hidden in boundary conditions. As they used an atmosphere-only GCM, boundary conditions (sea surface temperature and sea-ice temperature) had to be prescribed. They had been fixed to the present-day cycle. By performing experiments in which the prescribed sea-ice temperature was replaced by a simple sea-ice model, they showed that the ‘hidden calendar’ contributed to a bias in the results. As climate models of the current generation typically contains at least simple ocean and sea-ice models, the problem of a calendar hidden in the boundary conditions is reduced, but can still occur due to other parameters, e.g. related to vegetation. The magnitude of this ‘calendar problem’ varies in time and is mainly related to the climatic precession. To illustrate the effect, Fig. 2.3 shows the day in the year (counted from 1 January) of the equinox in boreal autumn (true solar longitude is equal to 180 , with vernal equinox fixed at 21 March). The largest deviation for autumnal equinox during the last 500 000 years occurred at 198 kyr BP (8.9.: 15 days) and 209 kyr BP (1.10.: þ13 days). The difference between these two extremes is therefore 29 days. During this period, eccentricity had its Day of autumnal equinox within the year

between the two definitions are visible on those latitudes, where absolute values show strong differences within a few days. This is mainly the case for high northern and southern latitudes (see Fig. 2.2a). On these latitudes, anomalies of more than 60 Wm2 occur when they are calculated on the basis of the classical calendar definition. If anomalies are calculated according to the astronomical longitude, they are close to zero for September and October in the high northern and southern latitudes.

21

275

270

265

260

255

250 0

–50 –100 –150 –200 –250 –300 –350 –400 –450 –500

Time (kyr) from 1950 AD

Fig. 2.3 Day at which the autumnal equinox occurs (i.e. true longitude of the Sun is 180 ), when vernal equinox is fixed at 21 March. Days are counted from 1 January.

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maximum. Owing to this variation in time, the consequences also have to be considered when interpreting transient simulations (see also discussion in Kubatzki et al., Chapter 39, this volume). 2.11 CHARACTERISTICS OF INSOLATION DURING INTERGLACIALS In this section, we will discuss the insolation during the interglacials, with a special emphasis on the interglacial prior to the Holocene, i.e. marine isotope substage (MIS) 5e (also referred to as 5.5.) in the marine records, more or less equivalent to the Eemian in the continental records. It is characterized by large values of the eccentricity (up to 0.0414 at 115 kyr BP), although the largest values of the last 500 kyr are reached during MIS 7 (0.0503 at 207 kyr BP). As eccentricity modulates the amplitude of the precession signal, this is also large. Moreover, according to climatic precession, NH summer occurred at perihelion at 127 kyr BP and at aphelion at 116 kyr BP. Over the same time, obliquity varies from a maximum value (24.259 ) at 131 kyr BP to a minimum value (22.316 ) at 112 kyr BP.

Consequently, the insolation at 65 N at the summer (June) solstice is maximum at 128 kyr BP ð547 Wm2 Þ. It is actually the largest value reached over the last 200 000 yr. Nevertheless, it must be kept in mind that the climate system is driven by the time evolution of the distribution of insolation during the year and along the latitude, so other latitudes and days deserve to be accounted for (see Fig. 2.2a to 2.2f). Some features of the interglacials of the last 500 kyr are summarized in Table 2.1. Figure 2.2e, 2.2c and 2.2f show the distribution of insolation over latitudes and seasons for respectively 128 kyr BP, 125 kyr BP and 115 kyr BP as anomaly from today. 125 kyr BP is often used for equilibrium experiments with GCMs (e.g. chapters 33, 34, 38). Although 3000 years after the maximum of insolation at 65 N, this date is selected because conventionally the minimum global ice cover occurs approximately at that time, which is considered as the climate optimum (Kukla et al., 2002). For example, insolation at 65 N decreases by 2.4% between 128 and 125 kyr BP. Compared to these dates, the angle of perihelion at 115 kyr BP was almost opposite, but similar to today. This date represents approximately the end of the warm phase (Kukla

Table 2.1 Orbital features and mid-month June insolation at selected dates during interglacials of the last 500 000 years

MIS 1 MIS 5 MIS 7 MIS 9 MIS 11

Date1 (ka)

Eccentricity

Climatic precession2

Obliquity (degrees)

Mid-month June insolation at 65 N (Wm2 )

6 11 122 128 216 219 330 334 405 411

0.018682 0.019529 0.040744 0.039017 0.048858 0.047839 0.034387 0.031539 0.019535 0.019057

FE SS FE SS FE (end summer) SS (early summer) FE (end summer) SS FE (end summer) SS (early summer)

24.105 24.201 23.336 24.130 24.308 23.947 24.099 24.238 23.208 24.023

506.60 528.45 498.23 548.32 529.86 551.96 514.61 541.95 496.22 522.68

1 For each marine isotope stage (MIS), the first date corresponds to the peak of the interglacial according to SPECMAP (Martinson et al., 1987) and the second one to the maximum of mid-month June insolation prior to the peak of the interglacial. 2 Here we give the approximate time in the year at perihelion (FE stands for fall equinox and SS for summer solstice). The exact time of passage at perihelion can be obtained from the values in the supplementary tables.


et al., 2002) and is therefore used in experiments examining the start of the glaciation. It must be kept in mind however that it is difficult to have an absolute dating of palaeoclimate records. At best, the radioactive methods are accurate within a few per cent. For stage 5, it means a few thousands years which is therefore more or less the same order of magnitude as the assumed lag between insolation and proxy records. Let us first consider the long-term variations of the deviation from present day of the insolation at the summer (June) solstice for time slices of 30 kyr centred at the peak of the interglacials from the northern to the southern poles (Fig. 2.4 and figures in the supplementary material). This latitudinal distribution around MIS 5e (Fig. 2.4) displays strong positive deviations going from 137 to 121 kyr BP, over all the latitudes. The largest deviation is in the northern polar regions at 128 kyr BP (larger than 70 Wm2 ). Negative deviations follow from 121 to 111 kyr, with the most negative value at the North Pole at 116 kyr BP (lower than 40 Wm2 ). Between these dates, the rate of change at the North Pole is about 10 Wm2

23

per kyr. It is of the same order for the other interglacials, except for MIS 11 and MIS 1 where it is slightly lower. This pattern of large positive deviations before the peak of the interglacial, much smaller deviations at the peak of the interglacial (even negative deviations) and large negative deviations after the peak of the interglacial is similar for the different interglacials over the last 500 kyr (supplementary figures). However, the most negative deviation after MIS 7.3 does not appear in the polar region but rather in the equatorial one. The peak of the interglacials lags behind the insolation maximum by 3 to 6 kyr. To follow, the deviations from present day of the insolation at 65 N will be investigated all through the year for the same 30-kyr time intervals (Fig. 2.5 and supplemental material). Over the time interval 137–107 kyr BP, the maximum positive deviation occurs in June at 128 kyr BP (Fig. 2.5). Between 137 and 128 kyr BP, the largest deviation occurs earlier in the year; it occurs later in the year after 128 kyr (Fig. 2.5). This is related to the 2 kyr phase shift of the insolation, from one month to the next. Moreover, the

80N –40

60N –30

40N

60

Latitude

20N –20

EQ 40 40

–10

20S

30

0

40S 60S

20 10

20

10

80S –108

–111

–114

–117

–120

–123

–126

–129

–132

–135

Time (kyr) from 1950 AD –70 –60 –50 –40 –35 –30 –25 –20 –15 –10 –5

5

10

15

20

25

30

35

40

50

60

70

Fig. 2.4 Latitudinal distribution of the time evolution of mid-month insolation (Wm2) at June solstice from 137 to 107 kyr (i.e. including MIS 5.5) as deviation from the present-day values.

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A. Berger et al.

–108

–10

–111 10

–114

Time (kyr) from 1950 AD

–20

20

–117 10

–120

20 –20 40

–123 –10

–126 –129 –10

60

–132 40

–135

20 10

–60 Jan

–30 Feb

0 (VE)

30 Apr

60 May

90 Jun

120 Jul

150 Aug

180 Sep

210 Oct

240 Nov

270 Dec

True longitude on earth’s orbit from vernal equinox (VE) –80 –70 –60 –50 –40 –30 –25 –20 –15 –10 –5

5

10 15 20 25

30 40 50 60 70 80

Fig. 2.5 Time evolution of the annual cycle of mid-month insolation at 65 N (Wm2) from 137 to 107 kyr as deviation from the present-day value.

following largest negative deviation after the peak of the interglacial also occurs in June. The general pattern is similar for the other interglacials (supplementary figures), although the amplitude is smaller then, except for MIS 7. The positive deviation is larger at MIS 7.3 than at MIS 5e. Moreover, the negative deviation at MIS 7.2 does not display one single maximum in June but rather a double one, in April and August. Finally, the long-term changes in the latitudinal and seasonal distributions of the deviations from present-day insolations can be documented over these time slices (Fig. 2.2a to 2.2f, supplementary figures and movies). For MIS 5e, starting from 137 kyr BP, the deviations become positive at the beginning of the year. The largest values are initially in the January–February SH and propagate very quickly (132 kyr BP) to the NH in the NH spring. The maximum deviation is then reached (128 kyr BP) in June at the North Pole. At that time, the deviations are positive in NH spring and NH summer over the whole Earth. Then, the positive deviation moves towards NH winter and it fades away. At the same time

(starting from 128 kyr BP), a large negative deviation appears, first in the tropical SH in January. At 122 kyr BP (the peak of the interglacial according to SPECMAP; Martinson et al., 1987), the deviation from the present day is negative in NH winter and NH spring, with a maximum negative value in the southern tropical area in February. It is positive in NH summer and NH autumn, with a maximum value in the northern tropical area in August. The negative deviations propagate towards the North with smaller amplitude and then culminate in the northern polar regions in June (116 kyr BP). The other interglacials display a similar behaviour, although the amplitude might be different. As discussed theoretically, the seasonal and annual irradiations (Fig. 2.6) at a given latitude only depend on the obliquity (as well as very slightly on eccentricity). Moreover, the length of the season only depends upon precession. Consequently, the seasonal mean irradiance is a function of both obliquity and precession. For MIS 5e, the total irradiation during the astronomical NH summer at 65 N displays a maximum


maximum at 127 kyr BP. It must be underlined that, for all the interglacials, the timing of maximum of NH summer mean irradiance is driven by the timing of the minimum of the length of the season and not by the timing of the total seasonal irradiation. Moreover, the interglacial peaks only a few thousand years after the maximum in NH summer mean irradiance. The strong imprint of obliquity on the total NH summer and annual irradiations also shows up very clearly in the latitudinal distribution of their deviations from the present day (Fig. 2.7 and supplementary material). For the NH summer, there is a latitudinal reversal in the sign of the deviation at 12 N. This reversal is double, at 42 N and 42 S, for the annual value, that is, the deviation has the same sign polewards of these latitudes. During the last interglacial, a reversal in time in the total NH summer irradiation occurs at 122 kyr BP, that is, at the peak of the interglacial. However, it is not a general feature of all the interglacials. Indeed, the peak of the interglacial occurs either during a positive deviation in the NH (MIS 1, MIS 7 and

BER78

3000 2950 2900 2850

100

2800

95 90 85

420 400 380 360 340

215

210 0

–100

–200

–300

–400

Fig. 2.6 From top to bottom: time evolution between 450 and 50 kyr AP of the total summer irradiation (106Jm2) at 65 N, the length of the summer season (day), the summer mean irradiance (Wm2) at 65 N and the annual mean irradiance (Wm2) at 65 N.

value at 131 kyr BP, NH summer is the shortest at 127 kyr BP, and consequently, NH summer mean irradiance reaches a 80N

25

–140

100

–100

60N

60 –60

40N 20

–20

Latitude

20N EQ

20 –20

20S 40S 60S –20

20

80S –108

–111

–114

–117

–120

–123

–126

–129

–132

–135

Time (kyr) from 1950 AD –160 –140 –120 –100 –80 –60 –40 –20

–5

0

5

20

40

60

80

100 120 140 160

Fig. 2.7 Variation between 137 and 107 kyr BP of the deviation from the present-day value of the summer irradiation (JJA, 106Jm2).

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A. Berger et al.

80N 60N

Latitude

40N 20N

–20

EQ

40 –10

20S

–10

20

40S

20 0

60S

10

10

80S –108

–111

–114

–117

–120

–123

–126

–129

–132

–135

Time (kyr) from 1950 AD –60

–50

–40

–30

–20

–10

–5

0

5

10

20

30

40

50

60

Fig. 2.8 Variation between 137 and 107 kyr BP of the deviation from the present-day value of the summer mean irradiance (JJA, Wm2).

MIS 9), during a negative deviation in the Northern Hemisphere (MIS 11) or at the transition (MIS 5). There is a better agreement between the pattern of the deviation for the JJA mean irradiance for different interglacials (Fig. 2.8 and supplementary figures). The positive deviation over the Northern Hemisphere prior to the peak of the interglacial is replaced by a negative deviation following it. The largest amplitudes are during MIS 7.3.

REFERENCES Adhe´mar, J.A., 1842. Re´volutions de la Mer: De´luges Pe´riodiques, Carilian-Goeury et V. Dalmont, Paris. Berger, A., 1978. Long-term variations of daily insolation and Quaternary climatic changes. Journal of Atmospheric Sciences, 35(12), 2362–2367. Berger, A., 1988. Milankovitch theory and climate. Reviews of Geophysics 26(4), 624–657. Berger, A., Loutre, M.F., 1994. Long-term variations of the astronomical seasons. In: Topics in Atmospheric and Interstellar Physics and Chemistry, Cl. Boutron (ed.), 33–61, Les Editions de Physique, Les Ulis, France.

Berger, A., Loutre, M.F., 1997. Intertropical latitudes and precessional and half precessional cycles. Science, 278, 1476–1478. Berger, A., Loutre, M-F., Tricot, C., 1993. Insolation and Earth’s orbital periods. Journal of Geophysical Research, 98 (D6), 10.341–10.362. Berger, A., Loutre, M.F., Me´lice, J.L., 2004. 100-kyr and 5.5-kyr periods in tropical insolation. AGU, PP16-A Tropical Perspectives on the Ice Ages, section Paleoceanography and Paleoclimatology, San Francisco, 14 December 2004. Brouwer, D., van Woerkom, A.J.J., 1950. Secular variations of the orbital elements of principal planets. Astron. Papers Am. Ephem., 13(2), 81–107. ¨ ber Bru¨ckner, E., Ko¨ppen, W., Wegener, A., 1925. U die Klimate der geologischen Vorzeit, Zeitschrift fu¨r Gletscherkunde, 14. Croll, J., 1875. Climate and Time in Their Geological Relations. Appleton, New York. Emiliani C.R.W., 1955. Pleistocene temperatures. Journal of Geology, 63(6), 538–578. Fro¨hlich, C., Lean, J., 2004. Solar radiative output and its variability: evidence and mechanisms. Astron. Astrophys. Rev., 12, 273–320. Hays, J.D., Imbrie, J., Shackleton, N.J., 1976. Variations in the Earth’s orbit : Pacemaker of the ice ages. Science, 194, 1121–1132. Joussaume, S., Braconnot, P., 1997. Sensitivity of paleoclimate simulation results to season definition. Journal of Geophysical Research, 102(D2), 1943–1956.

Insolation During Interglacial Ko¨ppen V., Wegener, A., 1924. Die Klimate der geologischen Vorzeit. Verlag Gebru¨der Borntraeger, Berlin, 255 pp. Kukla, G.J., Bender, M.L., de Beaulieu, J.L., Bond, G., Broecker, W.S., Cleveringa, P.J., Gavin, J.E., Herbert, T.E., Imbrie, J., Jouzel, J., Keigwin, L.D., Knudsen, K.-L., McManus, J.F., Merkt, J., Muhs, D.R., Mu¨ller, H., Poore, R.Z., Porter, S.C., Seret, G., Shackleton, N.J., Turner, C., Tzedakis, P.C., Winograd, I.J., 2002. Last interglacial climates. Quaternary Research, 58, 2–13. Martinson, D.G., Pisias, N.G., Hayes, J.D., Imbrie, J., Moore, T.C., Shackleton, N.J., 1987. Age dating and the orbital theory of the ice ages: development of a high-resolution 0 to 300,000-year chronostratigraphy. Quaternary Research 27, 1–29. Milankovitch, M., 1941. Kanon der Erdbestrahlung und seine Anwendung auf das Eiszeitenproblem. Royal Serbian Sciences, Spec.pub.132, Section of Mathematical and Natural Sciences, Vol. 33, pp. 633, Belgrade (‘‘Canon of Insolation and the Ice Age problem’’, English Translation by Israe¨l Program for Scientific Translation and published for the U.S. Department of Commerce and the

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National Science Foundation, Washington D.C., 1969, and by Zavod za Udzbenike I nastavna Sredstva in cooperation with muzej nauke I technike Srpske akademije nauka I umetnosti, Beograd, 1998). Murphy, J.J., 1876. The glacial climate and the polar ice-cap. Q.J. Geol. Soc. London, 32, 400–406. Paillard, D., 2001. Glacial cycles: Toward a new paradigm. Reviews of Geophysics, 39(3), 325–346. Penck, A., Bru¨ckner, E., 1909. Die Alpen im Eiszeitalter, Tauchnitz, Leipzig. Solanki, S.K., Usoskin, I.G., Kromer, B., Schu¨ssler, M., Beer, J., 2004. Unusual activity of the Sun during recent decades compared to the previous 11,000 year. Nature, 431, 1084–1087. Tricot, Ch., Berger, A., 1988. Sensitivity of present day climate to astronomical forcing. In: Long and Short Term Variability of Climate, H. Wanner, U. Siegenthaler (eds), 132–152, Earth Science Series, Springer Verlag. Vernekar, A.D., 1972. Long-period global variations of incoming solar radiation. Meteorol. Monograph, 12(34), 130 p.


3. A Survey of Hypotheses for the 100-kyr Cycle Martin Claussen1, Andre´ Berger2 and Hermann Held3 1

Meteorological Institute, University Hamburg, and Max Planck Institute for Meteorology, Bundesstr. 53, D-20146 Hamburg, Germany 2 Universite´ catholique de Louvain, Institut d’Astronomie et de Geóphysique G. Lemaıˆtre, 1348 Louvain-la-Neuve, Belgium 3 Potsdam Institute for Climate Impact Research, Telegrafenberg A31, D-14473 Potsdam, Germany

ABSTRACT Theories and mathematical models of longterm Quaternary climate variations are briefly summarized and revisited. We conclude that the problem of Quaternary climate variations, in particular the existence of a dominant 100-kyr ice-age cycle is, even approximately 160 years after first geological evidence of ice ages was found, not yet solved. However, we have some clues on what elements a theory of Quaternary Earth system dynamics should consist of. Assessment of a number of conceptual models – ranging from models in which forcing is necessary to yield observed climate variability to models of free climate oscillations – cannot favour any model over its competitor on the grounds of tuning each model to the time series of global ice volume. Hence, geographically explicit fully coupled climate system, or natural Earth system, models are required to analyse the system’s response to geographically varying forcing and internal feedbacks. Evidence emerges that much of Quaternary climate variability arises due to internal feedbacks, with ice sheets and biogeochemical cycles as critical elements and orbital forcing as pacemaker. 3.1 A BRIEF HISTORY OF THEORIES OF ICE AGES In 1842, briefly after L. Agassiz proposed the existence of ice ages on the grounds of

geological evidence, J.A. Adhe´mar suggested the first astronomical theory of climate change on the basis of the known precession of equinoxes. During the following decades, glacial geology became strongly tied to the astronomical theory which was advanced by J. Croll in the 1860s (references to Agassiz, Adhe´mar and Croll, see Paillard, 2001). Since Croll’s theory appeared to be more and more at variance with emerging geological evidence, his theory was eventually refuted. In 1896, Arrhenius concluded that ‘it seems that the great advantage which Croll’s hypothesis promised to geologists, viz. of giving them a natural chronology, predisposed them in favour of its acceptance. But this circumstance, which at first appeared advantageous, seems with the advance of investigation rather to militate against the theory, because it becomes more and more impossible to reconcile the chronology demanded by Croll’s hypothesis with the facts of observation’ (Arrhenius 1896, p. 274). The astronomical theory was modified and advanced by Milankovitch (1941) and by Ko¨ppen and Wegener (1924) at the beginning of the last century; however, it was disputed because it did not seem to be supported by geological data (Crowley and North, 1991). The orbital theory saw a strong revival after new geological evidence presented by J.D. Hays, J. Imbrie and N.J. Shackleton in 1976 (Hays et al., 1976) corroborated many of its predictions advanced and refined by A. Berger in the 1970s (Berger, 1977, 1978).

Martin Claussen, Andre´ Berger and Hermann Held

The theory of ice ages in which geochemical reactions and CO2 play a major role is perhaps as old as the astronomical theory. Early work goes back to J.J. Ebelmen in the 1840s, J. Tyndall in 1861 and S. Arrhenius in 1896. Arrhenius, for example, was convinced that changes in atmospheric transparency (due to changes in atmospheric CO2) would ‘prove useful in explaining some points in geological climatology which have hitherto proved most difficult to interpret’ (Arrhenius 1896, p. 275). Also today, there are models, e.g. by G. Shaffer and M.E. Raymo, developed in the 1990s (Shaffer, 1990; Raymo et al., 1997 ), which could be described as biogeochemical oscillators in which ocean biogeochemistry is the key player. 3.2 THE 100-kyr PARADOX Climate archives of the Pleistocene reveal variations in oxygen isotopes, with a dominant periodicity of approximately 40 kyr in the early Pleistocene, approximately 2 to 1 million years before present. In the late Pleistocene, the dominant periodicity of climate variations shifts to 100 kyr (e.g. Ruddiman et al., 1986). The amplitude of climate variations – interpreted as change in temperature and ice volume – seemingly increased. In particular, a tendency towards warmer, or more ice free, interglacials was found for the last 500 kyr in ice cores (see EPICA community Members, 2004, and figures in Chapter 4) as well as in marine isotopes (Fig. 3.1). At a first glance, the periodicities found in proxy data seem to coincide with those of the three astronomical parameters which characterize the Milankovitch theory. Milankovitch (1941) showed that meridional and seasonal changes in insolation consist of changes in the eccentricity of the Earth’s orbit, in the obliquity and in the precession of equinoxes. Berger (1977, 1978) demonstrated from analytical

–0.5

δ18O (‰)

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–1.0 –1.5 –2.0 2.0

1.5

1.0

0.5

0.0

Age (million years)

Fig. 3.1 Changes in 18 O reconstructed from marine sediment cores for the last 2 million years. These changes are interpreted as changes in global ice volume where increasing 18 O-values indicate increasing ice volume. The curve represents data from the tropical western Pacific sedimentary core ODP 806. This figure is taken from Saltzmann (2002) with permission of Academic Press.

developments of these astronomical parameters that the dominant periods are 400, 125 and 95 kyr for eccentricity, a dominant period at 41 kyr for obliquity and a bimodal period at 23 and 19 kyr for precession. However, the amplitudes of insolation changes due to changes in eccentricity appear to be much weaker than the amplitudes associated with obliquity and precessional variations, while proxy data show just the opposite. Hence, all theories employing orbital forcing in some way or the other need to resolve this so-called ‘100-kyr paradox’: the coexistence of a weak 100-kyr component in the astronomical forcing and a strong one in the response. The paradox also applies to theories involving other astronomical parameters than those used in the traditional Milankovitch theory. Again, the radiative forcing implied by these other astronomical forcings appears to be rather small, and thus, some additional amplifier is required to explain the dominance of the 100-kyr period. 3.3 FREE MODELS A wealth of models have been suggested to explain Pleistocene climate variations, in particular the 100-kyr cycle. According to Saltzman (2002), these models can be categorized into ‘forced’ and ‘free’ models. In

A Survey of Hypotheses for the 100-kyr Cycle

forced models, astronomical forcing (which explicitly includes the frequencies of orbital variations) is necessary to obtain the observed frequencies of climate variations. Free models do not rely on astronomical forcing at all, but the 100-kyr period appears as an internal (free) oscillation perhaps due to an internal instability of the system. In terms of dynamical systems analysis, the 100-kyr oscillation is the consequence of a bifurcation* to a self-sustained oscillation which is driven by an instationarity of the Earth system. Such free oscillators can address ice masses interacting with another variable: either ice-sheet location, bedrock depression or the thermohaline circulation. The free oscillator is forced by steady long-term variations, for example by a tectonic forcing associated with a slow variation in CO2 outgasing. An illustrative example of a free model is given by Paul and Berger (1997). They construct a nonlinear oscillator, which mimics ice-sheet dynamics as function of insolation, accumulation and bedrock response:

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glacial – interglacial sea-level variations of the late Pleistocene when orbital forcing is applied – of course, because it is designed to do so. A conceptual model like this is not designed as a predictive tool, but an exploratory tool. An interesting result, however, is that even with a random, white noise forcing, the model recaptures the asymmetric, sawtooth like 100-kyr cycle. This result suggests that the 100-kyr cycle could emerge as an internal property of the dyna-mics of some climate system component – for example the ice sheet – independent of any specific cyclicity in the forcing. Another example of a (more elaborate) freely oscillating model that includes ice-sheet dynamics with basal melting and sliding is given by Saltzman and Verbitzky (1993). An essential point is that the free models do not require amplification of small forcing and explain climate changes over the Pleistocene as simply the response of an internal oscillator that could in some cases even resist additional forcing over a variety of timescales.

dV ¼ a þ ib V c ðmðVÞÞd dt where t is the time, V is the nondimensional ice volume, a is a constant accumulation, i is the normalized ð0 < i < 1Þ insolation and m is a memory function m ¼ Rt m0 þ 1=T tT Vðt9 Þdt9 which could be interpreted as bedrock response. The forcing i(t) is chosen as summer insolation, either monthly mean or summer mean or annual maximum insolation at 65 N. That boreal summer insolation could the decisive forcing parameter of Pleistocene ice ages was suggested by Wladimir Ko¨ppen based on suggestions by Albrecht Penck and Eduard Bru¨ckner (see Ko¨ppen and Wegener, 1924; Milankovitch, 1941). The unknown parameters a, b, c, d, m0 and T are tuned to data. This model perfectly reproduces the *

(i.e. a qualitative change of the system’s state caused by a change in a control parameter, may it be a key, internally generated slow variable or an external forcing).

3.4 FORCED MODELS Saltzman (2002) provides a table in which he summarizes the various types of forced models. Here we would like to highlight only a few types. The simplest forced models are linear models of global ice mass in which ice-sheet physics and feedback processes with the atmosphere or surficial physics are neglected. For example, Imbrie (1980, see Paillard, 2001) explores the consequences of a simple low-pass filter that relates the dynamic change of nondimensional ice volume V to insolation i by dV ði VÞ ¼ dt where is a response time which differs for ablation (i.e. when dV=dt < 0) and

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accumulation (when dV=dt 0). Imbrie’s model underestimates the 100-kyr cycle. Similar is valid for the simple response/ threshold model by Calder (1974, see Paillard, 2001) in which dV ¼ kði i0 Þ dt where i0 is a threshold of insolation and k is a parameter which differs for accumulation and ablation. Calder’s idea of a threshold-dominated system was more consequently formulated by Paillard (1998). Paillard assumes that the climate system exhibits multiple equilibria. If a certain threshold in the forcing is reached, then the system would jump from one equilibrium to another. In Paillard’s model, these equilibria are associated with an interglacial state, a glacial maximum and a weak glacial, such as the early Weichselian, for example. Again, the parameters (in this case the threshold values in insolation in Paillard’s model) are tuned to data. Paillard’s model provides a new view on the climate system as a threshold system. In particular, it solves the paradox of long interglacials occurring at seemingly low insolation: it suggests that not the insolation itself is the important driver, but the variation of insolation which would cause certain thresholds to be crossed more or less often. During long interglacials, like that at MIS (marine isotope stage) 11, insolation is much smaller than compared to the Eemian interglacial, but, because of low eccentricity, it varies only marginally for a longer period of time. Threshold models can become even more sophisticated, if in addition, stochasticity is present. Then the periodic forcing may be less pronounced and even too weak to cause transitions between equilibria in the noisefree case, and yet, with noise being added, it can induce synchronized jumps due to stochastic resonance (e.g. Benzi et al., 1982). Again, it is important for this theory to work that the system reveals multiple equilibria.

Hence, the external driver is not really a forcing, but rather a trigger. In general, stochastic climate models explore the consequences of random walk processes. This idea was originally proposed by Hasselmann (1976). Hasselmann assumed that the annual cycle of insolation generates variability in the fast climate components which could be randomly accumulated by the more sluggish climate components. Recently, Wunsch (2003) pursued this idea by demonstrating that most long-term climate archives reveal a rednoise spectrum, that is, a spectrum with amplitude of variance decaying with larger frequencies. Superimposed on the red spectrum are weak structures corresponding to the frequency bands of an orbital forcing. To explain the dominant 100-kyr climate variability, Wunsch suggests a stochastic forcing of a system with a collapse threshold. In such a system, a timescale can be generated without having to assume an external frequency component or an internal resonance, but the average time of a Brownian random walk to reach the threshold is sufficient to explain an average period of 100 kyr. Furthermore, Wunsch’s model yields a transition in the spectral domain from red to white, that is, a flat spectrum with amplitudes independent of frequency. This way, variability on the 20–100-kyr timescale, as well as on shorter timescales can be explained in a combined way. The categorization of models of Quaternary climate system dynamics as ‘forced’ and ‘free’ can be complemented by ordering them by means of complexity in terms of physical processes involved. According to Saltzman (2002, p. 276), climate system models aim at ‘ever more complete representation of the full slow-response climate system’, hence the centre manifold.* *

In the ‘vicinity’ of a bifurcation (see above), one can identify slow and fast processes. Then, the centre manifold characterizes the interplay of the few slowest variables with the collective effect of all the other variables, and therefore allows to study a complex system’s long-time behaviour by just a few degrees of freedom, hence by a conceptual model.

A Survey of Hypotheses for the 100-kyr Cycle

Saltzman (2002, p. 276) suggests that ‘ice sheets and their bedrock and basal properties, coupled with forced and free variations of carbon dioxide, operating on an Earth characterized by a high-inertia deep thermohaline ocean that can store carbon and heat . . .’ encompass the centre manifold. 3.5 WHAT TYPE OF MODEL DO WE NEED? So far, only conceptual, or inductive, models have been discussed. These models are designed to demonstrate the plausibility of processes; they are based on a gross understanding of feedbacks that are likely to be involved (Saltzman, 1985). Roe and Allen (1999) investigated the performance of six representatives of inductive models. They tuned each model to the time series (the last 900 kyr) of global ice volume and also the time rate of change for the ice volume while modelling the residuals as first- (and second) order autoregressive process. They found that within 95% error bars, one cannot favour any model over its competitors. Obviously, some assumptions on which conceptual models are based on are oversimplified and thus, could even be misleading. For example, the assumption that orbital forcing is identified with summer insolation at high northern latitudes overemphasizes the ice albedo feedback at high northern latitudes and neglects the fact that other components of the climate system can react to varying insolation in a different way than Northern Hemisphere ice sheets do. Therefore, it seems sensible to explore the role of geographically varying forcing and feedback processes in geographically explicit models. The degree of spatial and temporal resolution necessary for palaeoclimate simulations is disputed. Comprehensive, ‘state-of-the-art’ coupled models describing the general circulation of the atmosphere and the ocean, the dynamics of the terrestrial biosphere and the ice sheet as well as

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biogeochemistry are supposed to be the most realistic laboratory of the natural Earth system. However, their applicability to long-term simulations is limited by high computational costs. Therefore it was proposed to use Earth system models of intermediate complexity (EMICs; Claussen et al., 2002) which operate at a higher level of spatial and temporal aggregation. Many processes resolved in AOGCMs have to be parameterized in EMICs. The advantage gained by this reduction is computational efficiency which makes them a useful tool for integrated palaeoclimate modelling. There are already numerous palaeoclimatic studies using EMICs and AOGCMs, and some examples will be given in Section 5 of this book. So far, however, even with EMICs, a realistic simulation of glacial – interglacial cycles by using orbital forcing only remains difficult to be done. 3.6 PERSPECTIVE Even approximately 160 years after first geological evidence, the ice-age riddle is not yet fully solved. However, we have some clues on what elements a theory of Quaternary Earth system dynamics should consist of – regarding concepts and model structure. Saltzman (2002) has proposed a unified theory of Quaternary Earth system dynamics. The term ‘unified’ is used, because it combines theories based on orbital forcing and on greenhouse gas forcing, respectively. Saltzman supposes that the slow part of the climate system involves the ice sheets and their bedrock and basal properties, coupled with tectonically forced and free variations of carbon dioxide, and a high-inertia deep thermohaline ocean. While it is very likely that slow variables exist, it is still not obvious that those can be identified with particular physical entities just mentioned. Quite the contrary, in a highly resolved spatiotemporal dynamics, such slow variables may emerge as complex patterns across physical entities,

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which strongly support the use of spatially resolved climate models. Validation of inductive models appears to be an almost futile task: assessment of a number of inductive models cannot favour any model over its competitor on the grounds of tuning each model to the time series of global ice volume. In the range of more comprehensive, quasi-deductive models, only EMICs have been used for long-term studies. Some of these numerical experiments (e.g. Galleé et al., 1992; Calov et al., 2005) support the idea of Hays et al. (1976) that orbital forcing may act as a pacemaker of glacial–interglacial cycles. However, the situation is rather complex. Obviously the response of the Earth system to a given forcing is a function of the actual state of the Earth system as well as meridional and seasonal changes of the forcing. Presumably, changes in insolation associated with changes in orbital parameters trigger fast internal feedbacks such as the water vapour – temperature feedback and the snow – albedo feedback which then are further amplified by slower feedbacks such as biogeochemical and biogeophysical feedback and the isostatic response of the lithosphere to ice-sheet loading. Some of these feedbacks even change sign during the course of a glacial–interglacial cycle. In Chapter 4, an example of exploring the feedbacks involved in the dynamics of glacial–interglacial climate change will be presented, while in chapters 5.1 to 5.8, the discussion is focused on the forcing and feedbacks at the end of an interglacial.

REFERENCES Arrhenius, S., 1896. On the influence of carbonic acid in the air upon the temperature of the ground. Philosophical Magazine and Journal of Science, 41, 237–276. Benzi, R.A., Parisi, G., Sutera, A., Vulpiani, A., 1982. Stochastic resonance in climatic change. Tellus, 34, 10–16. Berger, A., 1977. Support for the astronomical theory of climatic change. Nature, 268, 44–45.

Berger, A., 1978. Long-term variations of daily insolation and Quaternary climatic changes. Journal of Atmospheric and Oceanic Science, 35, 2362–2367. Calov, R., Ganopolski, A., Petoukhov, V., Claussen, M., Greve, R., 2005. Transient simulation of the last glacial inception. Part I: Glacial inception as a bifurcation in the climate system. Climate Dynamics, 25(6), 545–562. Claussen, M., Mysak, L.A., Weaver, A.J., Crucifix, M., Fichefet, T., Loutre, M.-F., Weber, S.L., Alcamo, J., Alexeev, V.A., Berger, A., Calov, R., Ganopolski, A., Goosse, H., Lohman, G., Lunkeit, F., Mokhov, I.I., Petoukhov, V., Stone, P., Wang, Zh., 2002. Earth system models of intermediate complexity: Closing the gap in the spectrum of climate system models. Climate Dynamics, 18, 579–586. Crowley, T.J., North, G.R., 1991. Paleoclimatology. Oxford University Press, 339 pp. EPICA Community Members, 2004. Eight glacial cycles from an Antarctic ice core. Nature, 429, 623–628. Galleé, H., van Ypersele, J.-P., Fichefet, T., Marsiat, I., Tricot, C., Berger, A., 1992. Simulation of the last glacial cycle by a coupled, sectorially averaged climate-ice sheet model. Part II: Response to insolation and CO2 variation. Journal of Geophysical Research, 97, 15,713–15,740. Hasselmann, K., 1976. Stochastic models. I. Theory. Tellus, 28, 473–485. Hays, J.D., Imbrie, J., Shackleton, N.J., 1976. Variations in the earth’s orbit: pacemaker of the ice ages. Science, 194, 1121–1132. Ko¨ppen, W., Wegener, A., 1924. Die Klimate der geologischen Vorzeit. Borntraeger, Berlin, 255 pp. Milankovitch, M., 1941. Kanon der Erdbestrahlung und seine Anwendung auf das Eiszeitenproblem. Ko¨niglich Serbische Akademie, Belgrad, 633 pp. Paillard, D., 1998. The timing of Pleistocene glaciations from a simple multiple-state climate model. Nature, 391, 378–381. Paillard, D., 2001. Glacial cycles: towards a new paradigm. Review of Geophysics, 39, 325–346. Paul, A., Berger, W.H., 1997. Modellierung der Eiszeiten: Klimazyklen und Klimau¨berga¨nge. Geowissenschaften, 15, 20–27. Raymo, M.E., Oppo, D.W., Curry, W., 1997. The midPleistocene climate transition: a deep sea carbon isotopic perspective. Paleoceanography, 12, 546–559. Roe, G.H., Allen, M.R., 1999. A comparison of competing explanations for the 100,000-yr ice age cycle. Geophysical Research Letter, 26, 2259–2262. Ruddiman W.F., Raymo, M.E., McIntyre, A., 1986. Matuyama 41,000-year cycles: North Atlantic Ocean and Northern Hemisphere ice sheets. Earth and Planetary Science Letters, 80, 117–129.

A Survey of Hypotheses for the 100-kyr Cycle Saltzman, B., 1985. Paleoclimatic modeling. In: Hecht, A.D. (ed.), Paleoclimate analysis and modeling. Wiley, 341–396. Saltzman, B., 2002. Dynamical Paleclimatology: Generalized Theory of Global Climate Change. International Geophsics Series, Vol. 80, Academic Press, San Diego, 354 pp. Saltzman, B., Verbitzky, M.Y., 1993. Multiple instabilities and modes of glacial rhythmicity in the Plio-

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Pleistocene: a general theory of late Cenozoic climatic change. Climate Dynamics, 9, 1–15. Shaffer, G., 1990. A non-linear climate oscillator controlled by biogeochemical cycling in the ocean: an alternative model of Quaternary ice age cycles. Climate Dynamics, 4, 127–143. Wunsch, C., 2003. The spectral description of climate change including the 100ky energy. Climate Dynamics, 20, 353–363.


4. Modelling the 100-kyr Cycle – An Example From LLN EMICs Andre´ Berger and Marie-France Loutre Universite´ catholique de Louvain, Institut d’Astronomie et de Geóphysique G. Lemaıˆtre, 1348 Louvain-la-Neuve, Belgium

ABSTRACT The glacial–interglacial cycles have been reproduced by the LLN-2D Northern Hemisphere model over the whole Quaternary. A short description of the model is presented before the main results are presented for the last cycle, the last 200 kyr, the last 800 kyr and the transition between the 41-kyr and the 100-kyr world. The quality and the deficiencies of these results are discussed in relationship with proxy geological records originating from some deep-sea and ice cores. 4.1 INTRODUCTION The 100-kyr cycles in eccentricity result from combination of the periods associated with the first terms in the expansion of the climatic precession parameter (e.g. Berger, 1994). The signal of these 100-kyr cycles is very weak in the insolation spectra (Berger et al., 1993c). This is why, in the previous chapter, it was mentioned that, although the eccentricity cycle seems in phase with the 100-kyr cycle in most climatic records, it needs a nonlinear amplification by mechanisms such as those related to the ice sheets, the ice albedoand the water vapour–temperature feedbacks, the carbon cycle, the isostatic rebound, the deep-ocean circulation and/ or the ocean–ice interactions. This is the reason for which a model of the fully coupled climate system needs to be constructed to allow at least some of the interactions between the atmosphere, the hydrosphere, the cryosphere, the

biosphere and the lithosphere to be taken into account. Such a climate model was built in Louvain-la-Neuve in the late 1980s. It links the Northern Hemisphere atmosphere, ocean mixed layer, sea ice, ice sheets and continents (Galleé et al., 1991). It is a twodimension (latitude–altitude) sectorially averaged model. In each latitudinal belt, the surface is divided into at most seven oceanic or continental surface types, each of which interacts separately with the subsurface and the atmosphere. Special attention is paid to the albedo of snow, of vegetation in the northern high latitudes and of sea ice. The atmosphere–ocean model is asynchronously coupled to a model of the three main Northern Hemisphere ice sheets and their underlying bedrock. The coupled climate model is then forced by the astronomically derived insolation for each day and latitude and by the atmospheric CO2 concentration (this model did not contain an interactive carbon cycle). More details on the model are given in Galleé et al. (1991) and also in Berger et al. (1990) for the ice sheet–lithosphere model, in Berger et al. (1989) for the upper ocean and in Berger et al. (1994) for the radiative convective scheme. The model is able to reproduce the main features of the present-day atmospheric general circulation and seasonal cycles of the oceanic mixed layer, of the sea ice and of the snow cover (Galleé et al., 1991). This model has been further enlarged, leading to the MoBidiC model which has both the Northern and the Southern Hemispheres and a much more elaborated ocean (Crucifix et al., 2002).

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Andre´ Berger and Marie-France Loutre

4.2 LAST GLACIAL–INTERGLACIAL CYCLE A first set of modelling experiments (Galleé et al., 1991) showed that the variations in the Earth’s insolation alone induce feedbacks in the climate system which are sufficient to amplify the direct radiative impact and generate large climatic changes, provided CO2 is kept below 240 ppmv. This result confirms the (Hays et al., 1976) idea that the orbital forcing acts as a pacemaker of the ice ages, but initiation and termination of glacial cycles cannot be explained without invoking both the fast feedbacks associated with atmospheric processes (water vapour, cloud, snow and sea ice) and the slower feedbacks associated with other parts of the climate system, in particular build-up and disintegration of ice sheets. Moreover, taking into account the Vostok CO2 variations (Barnola et al., 1987) allows to better shape the last glacial–interglacial cycle and in particular the air temperature (Galleé et al., 1992; Loutre et al., 1994). In a similar experiment with MoBidiC, there is a generally increasing trend in the continental ice volume from the last interglacial until the last glacial maximum. This maximum in the Northern Hemisphere is reached at 18 kyr BP with a volume of more than 40 106 km3 , a value probably slightly underestimated when compared to empirical reconstructions and other simulations (Hagdorn, in preparation). The simulated annually averaged surface temperature varies by up to 2.1 C in global average. These variations are greater for the Northern Hemisphere (2.8 C) than for the Southern Hemisphere (1.5 C), but larger variations can occur in higher latitudes. In the 50–60 N latitude band, the annual mean surface temperature of the Eurasian continent increases by 3.7 C between 25 kyr BP and the present, with an amplitude of the variations much larger in monthly than in annual means. The high northern latitudes of the Atlantic Ocean exhibit abrupt temperature changes of large amplitude at

around 100 and 10 kyr BP. Both events are characterized by a rapid cooling of 4 to 5 C followed by a period of slow warming, although temperatures remain low. Climate then warms rapidly by 6–7 C. In both cases, the export of NADW (North Atlantic Deep Water) is reduced by 5 Sv (1 Sv ¼ 106 m3 s1 ). These events are also characterized by a southern shift of the convection zone in the Atlantic and a cooling of the upper and intermediate waters of the Atlantic Ocean. The total area and volume of Arctic sea ice also increase during these events. 4.3 LAST 200 kyr Because of the sensitivity of the nonlinear climate model, it was crucial to see whether it can sustain more than one glacial– interglacial cycle. This was at the origin of a second simulation covering the last 200 kyr (Galleé et al., 1993). Both the insolation forcing and the CO2 variations reconstructed from deep-sea (Shackleton et al., 1992) and ice (Jouzel et al., 1993) cores were used (Fig. 4.1 top). Broadly speaking, the response of the model to the Vostok CO2 and to the insolation (Berger, 1978) forcings reproduces quite well the low frequency part of the geological record over the last 200 kyr (Fig. 4.2 top) (Berger and Loutre, 1996; Berger et al., 1998). The timing of the stacked, smoothed oxygen isotope record of SPECMAP (Imbrie et al., 1984; Martinson et al., 1987) compares favourably with the simulation, although it might be argued that this result is biased by the astronomical tuning of the SPECMAP record. Moreover, there are discrepancies in the magnitude of the simulated ice volume. The largest one is probably the too large ice melting simulated by the model around 170 kyr BP and induced by the large values of insolation around 175 kyr BP, although the ice volume maximum at 182 kyr BP seems to be well captured by the model. Either it is a deficiency of the model or we have to look for a

Modelling the 100-kyr Cycle 0

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Fig. 4.2 Simulated ice volume of the Northern Hemisphere using the LLN 2-D NH climate model forced by insolation (Berger, 1978) and the different scenarios of atmospheric CO2 concentration of Fig. 4.1. Black curve is the ice volume of SPECMAP (Imbrie et al., 1984).

40


significant change in the Southern Hemisphere continental ice at that time. At the end of stage 6, the simulated glacial maximum occurs at 135 kyr BP, while the 18 O ice volume maximum occurs at 151 kyr, but remains large from 156 to 133 kyr BP. The model simulates very well the transitions between isotopic stages 6 and 5 and between the isotopic substages 5e (Eemian interglacial) and 5d. This is due mainly to the insolation changes, but it is reinforced by important changes in the CO2 concentration and all the feedbacks in the model including those related to land surface cover (Berger, 2001). Although the timing of what may correspond to isotopic substages 5c, 5b and 5a is well reproduced, the model simulates a total melting of the Northern Hemisphere ice sheets between 126 and 117 kyr BP, 100 and 97 kyr BP and 83 and 74 kyr BP. Although this is not realistic (the Greenland ice sheet having survived at least over the last two to three glacial–interglacial cycles Dansgaard et al. (1993)), it does not prevent the ice sheets to grow again, leading to a 100-kyr quasi-cyclicity similar to the one seen in the geological data. The entrance into stage 4, starting at 70 kyr BP, is not rapid enough when compared to geological record and especially to the SPECMAP reconstruction, a failure that can be solved by taking into account both the volumes of snow and ice. Then there is a brief reversal between 60 and 50 kyr BP, followed by a re-growth of the ice sheets and leading to the last glacial maximum. The maximum amount of ice reaches 47 106 km3 at 15 kyr BP. Finally, the model reproduces the deglaciation from 15 to 3 kyr BP. It simulates correctly the total disappearance of the Eurasian ice sheet around 7 kyr BP, followed by the melting of the North American one, only the Greenland ice sheet being left in the Northern Hemisphere with roughly 2:96 106 km3 of ice. Since 3 kyr BP, the simulated Greenland ice volume in the absence of human perturbations is increasing slightly, reaching today 3:07 106 km3 , which represents

about the actual present-day value. In parallel, our simulated natural climate is slightly cooling since the peak of the Holocene. These experiments confirm that variations in the Earth’s orbit and related insolation act as a pacemaker of the ice ages (Hays et al., 1976). CO2 variations help to shape the 100-kyr cycle and mainly improve the simulated surface air temperature (Galleé et al., 1992; Loutre et al., 1994). A deeper analysis of these experiments also confirms the importance of the processes governing the response of the modelled climate system to insolation and/or CO2 changes. These are fundamentally related to the albedo- and water vapour feedbacks (Berger et al., 1993b), to the taigatundra direct and indirect impacts on high latitudes surface albedo (Kubatzki and Claussen, 1998; Berger, 2001), to the altitude and continental effects on the precipitations over the ice sheets, to the lagging lithospheric response to the ice-sheet loading (Crucifix et al., 2001) and to the mechanical destabilization of the ice sheets through the rapid melting of their southern front as compared with the northern one (Berger et al., 1992; Berger et al., 1993a). 4.4 LAST 800 kyr To further test the capacity of the model to sustain the 100-kyr cycle over a long period of time, four CO2 scenarios were actually used for simulating the past 575-kyr (Fig. 4.1 middle for 440 kyr) (Loutre and Berger, 2003). The first one is based on a multiple regression between the deep-sea record from Ontong Java Plateau in the western equatorial Pacific and the ice core CO2 from Antarctica (Berger W. et al., 1996). The second one is generated from a regression between the Vostok CO2 concentration and the SPECMAP oxygen isotope values calculated over the last 218 kyr and extended over the past 575-kyr (Li et al., 1998). The same procedure was applied for the third one, using the low latitude stacked 18 O record from marine core MD 900963 of site 677 (Bassinot et al.,

Modelling the 100-kyr Cycle

1994) instead of the SPECMAP record. At last, the CO2 concentration reconstructed from Vostok over the last four glacial–interglacial cycles was used (Petit et al., 1999). As this Vostok record extends only to 414 kyr BP, the CO2 from Li et al. (1998) was used from 575 to 440 kyr BP, and a linear interpolation between 440 and 414 kyr BP ensured the transition towards the Vostok record. The broad features of the results obtained from these different CO2 scenarios are pretty well similar (Fig. 4.2 middle). The spectra of the simulated ice-sheet volume and of the oxygen isotope proxy record are highly coherent in the frequency bands associated with the periods of 100, 41, 23 and 19 kyr. In these experiments, as in previous ones, the albedo- and water vapour– temperature feedbacks play a significant role in amplifying the forcings. According to a simulation of the last glacial maximum in response to insolation and CO2 , the water vapour feedback would explain 40% of the cooling (Berger et al., 1993b). Such a change in the water vapour content of the atmosphere has also been simulated by Li et al. (1998) over the last 500 kyr, with a vertically integrated value over the whole Northern Hemisphere in July varying from a little bit more than 30 kg/m2 during glacial times to about 55 during the interglacials. For the land surface–climate interaction, when insolation and CO2 , decrease (insolation leading CO2), both the snow fields and tundra lead to an increase of the surface albedo creating a positive feedback, which is reinforced by the subsequent decrease in the water vapour content of the atmosphere (see also Loutre et al., this volume). These feedbacks lead finally to the building up of the continental ice sheets which in turn enter the feedback loops. The major difference between these experiments covering the last four glacial– interglacial cycles arises during isotopic stages 11 and 10. From 400 to 350 kyr BP, the ice volume simulated with Vostok CO2 concentrations remains lower than 5 106 km3 over the whole interval

41

(Fig. 4.2), while in the other experiments, it remains at its minimum for only 10 kyr. This feature of the experiment using Vostok CO2 concentrations, if confirmed, is similar to what might happen to our stage 1, even more if the future CO2 level is kept high over a sufficiently long period of time. From 330 kyr BP (peak of MIS 9 interglacial) to 250 kyr BP (glacial maximum in MIS 8), SPECMAP displays a regular trend towards glaciation, while the simulation displays stadials–interstadials with full interglacial conditions. From stage 5e onwards, there are no real significant differences between these experiments and the experiment discussed in the previous section. This can be expected from the rather good agreement between the different CO2 reconstructions, except during the early part of stage 5e. However, this difference in the CO2 series does not lead to any significant difference between the simulated ice volumes, contrary to what happened at stage 11 because the amplitude in the insolation variation is much larger during MIS 5e than during MIS 11. The similarity in the CO2 concentration of stage 5e and stage 11 (long lasting maximum) and the difference in their insolation stress gain the importance of better understanding the relatively long interglacials (Loutre et al., this volume) during which high values of CO2 are sustained, whether the ice volume remains rather low (stage 11) or not (stage 5e) during the whole stage. Several other attempts were made to provide CO2 values prior to 400 kyr BP, it means prior to the end of the Vostok ice core. The recent EPICA core is expected to solve this problem. Before its CO2 became available, a linear regression was built between Vostok CO2 (Petit et al., 1999) and EPICA deuterium (EPICA, 2004) over the last 414 kyr and used in order to extend the CO2 record to the lower end of the EPICA core (710 kyr BP, Fig. 4.1 bottom). The chronology of Vostok–CO2 includes corrections made by Barnola (personal communication), (Parrenin et al., 2004; Raynaud et al., 2005).

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In this regression, from 710 to 430 kyr BP, CO2 values vary between 200 and 250 ppmv, a range that is much less than in the Vostok record. For this reason, another scenario was created with essentially a decrease of the CO2 minima to 180 ppmv. In both scenarios, the model continues to simulate the 100-kyr cycle before 400 kyr BP. However, over these earlier times, the Northern Hemisphere ice sheets are still disappearing during interglacials, and their size during glacials is similar to what is simulated for MIS 8 and MIS 10 using the Vostok–CO2 record. Although for the ice volume, both the phase and amplitude compare pretty well with the reconstructions over the last 400 kyr, none of the simulations reproduces a reduction in the amplitude of the glacial–interglacial cycles before MIS 11 (Fig. 4.2 bottom) as it is seen in both deep-sea and EPICA records (EPICA, 2004), that is, cool interglacial–cold glacial. That might be related to processes lacking in our model and which play an important role before MIS 11 or to a Southern Hemisphere much cooler than the Northern one (we must recall that our model is only simulating NH), although evidence for this must still be found. 4.5 FROM 41- TO 100-kyr CYCLE A final test which needed to be performed is related to the transition between the 41-kyr and the 100-kyr world. Can this be simulated by the same model without any re-tuning and what might be a possible cause for it? To analyse this, a linearly concentration from decreasing CO2 320 ppmv at 3 Myr BP to 200 ppmv at the last glacial maximum was used as a scenario (Berger et al., 1999). Actually, during the late Pliocene, the simulated ice volumes are small and interglacials are long, while long glacials with a large amount of ice prevail during the late Pleistocene. The model is thus able to reverse from a late Pliocene/

early Pleistocene climate dominated by warm interglacials to a late Pleistocene cold climate where glacials prevail. At the transition, CO2 is sufficiently low (240 ppmv in our model) to allow the ice sheets to start developing fully. A few sensitivity tests confirmed the need of crossing such a CO2 threshold to generate the 100-kyr cycles. If the CO2 concentration is above the threshold, no glacial–interglacial cycle can be generated because no ice sheet can develop. The spectral analysis performed on the simulated ice volume between 2 and 1 Myr BP shows that the most important periodicity is related to obliquity. For the last 1 Myr, the global spectrum is characterized by the obliquity and precession frequencies, but it is the variance components near 100 kyr which dominate. This is confirmed in an evolutive spectrum from 1.5 Myr BP to present, where in addition one can see that the obliquity signal remains more or less constant and the periods in the precessional band start to strengthen around 1.3 Myr BP, especially the 23-kyr cycle which splits into the 23.7- and 22.4-kyr periods as it is the case in the precession expansion (Berger, 1977). As a conclusion, these model results stress the role of the CO2 concentration crossing some threshold value to allow the climate system and its feedback mechanisms to respond nonlinearly to the astronomical forcing and hence to create the 100-kyr cycle as a result of a beat between the main precessional frequencies. 4.6 CONCLUSIONS The LLN-2D model succeeds to reproduce the 41-kyr cycle of the early and middle Pleistocene, up to about 1 Myr BP and the progressive transition towards the 100-kyr cycle which dominates the spectrum of the last 400 kyr. The transition coincides with the crossing of a CO2 threshold of 240 ppmv. This value is indeed sufficiently low to allow ice sheets to start building up and to be further sustained. The same simulation

Modelling the 100-kyr Cycle

shows also that stages 11 and 1 cannot be interglacials if CO2 is low. After MIS 11, the model simulates pretty well the last 400 kyr when forced with the long-term variations of insolation and CO2. Sensitivity analyses stress the role of CO2 during times of very low eccentricity. As a consequence, an entrance into glaciation now can be simulated only if CO2 remains below 240 ppmv. Between 400 and 900 kyr BP, the model simulates the 100-kyr cycles with reduced amplitude, but is globally too warm, not allowing cool interglacials and cold glacials.

REFERENCES Barnola, J.M., Raynaud, D., Korotkevitch, V.S., Lorius, C., 1987. Vostok ice core: a 160,000 year record of atmospheric CO2. Nature 329(6138), 408–414. Bassinot, F.C., Labeyrie, L.D., Vincent, E., Quidelleur, S., Shackleton, N.J., Lancelot, L.Y., 1994. The astronomical theory of climate and the age of the Brunhes–Matuyama magnetic reversal, Earth and Planetary Science Letters 126(1–3), 91–108. Berger, A., 1977. Support for the astronomical theory of climatic change. Nature 268, 44–45. Berger, A., 1978. Long-term variations of daily insolation and Quaternary climatic changes, Journal Atmospheric Science 35(12), 2362–2367. Berger, A., 1994. Astronomical theory of palaeoclimates, in: Cl. Boutron (Ed.), Topics in Atmospheric and Interstellar Physics and Chemistry, Les Ulis, France, 411–452. Berger, A., 2001. The role of CO2 , sealevel and vegetation during the Milankovitch forced glacialinterglacial cycles, in: C.U.H.L. Bengtsson (Ed.), Geosphere–Biosphere Interactions and Climate, Cambridge University Press, New York, 119–146. Berger, A., Loutre, M.F., 1996. Modelling the climate response to astronomical and CO2 forcings, C.R. Acad. Sci. Paris, t. 323(se´rie IIa), 1–16. Berger, A., Fichefet, T., Galleé, H., Tricot, C., Marsiat, L., van Ypersele, J.-P., 1989. Astronomical forcing of the last glacial–interglacial cycle, in: P. Crutzen, J.C. Ge´rard and R. Zander (Eds.), Our Changing Atmosphere, Universite´ de Lie`ge, Institut d’Astrophysique, Cointe-Ougreé, 353–382. Berger, A., Fichefet, T., Galleé, H., Tricot, C., Marsiat, L., van Ypersele, J.-P., 1990. Physical interactions within a coupled climate model over the last glacial-interglacial cycle, Philosophical Transactions of the Royal Society of Edinburgh: Earth Sciences 81(4), 357–369.

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Berger, A., Fichefet, T., Galleé, H., Tricot, C., van Ypersele, J.-P., 1992. Entering the glaciation with a 2-D coupled climate model, Quaternary Science Reviews 11(4), 481–493. Berger, A., Galleé, H., Tricot, C., 1993a. Glaciation and deglaciation mechanisms in a coupled 2-D climate-ice sheet model, Journal of Glaciology 39(131), 45–49. Berger, A., Tricot, C., Galleé, H., Loutre, M.F., 1993b. Water-vapor, CO2 and insolation over the last glacial-interglacial cycles, Philosophical Transactions of the Royal Society of London Series BBiological Sciences 341(1297), 253–261. Berger, A.M., Loutre, M.F., Tricot, Ch., 1993c. Insolation and the Earths orbital periods, Journal of Geophysical Research-Atmosphere 98(D6), 10 341–10 362. Berger, A., Tricot, C., Galleé, H., Fichefet, T., Loutre, M.F., 1994. The last two glacial–interglacial cycles simulated by the LLN model, in: J.-C. Duplessy, M.-T. Spyridakis (Eds.), Long Term Climatic Variations, Data and Modelling 22, Springer, Berlin, 411–452. Berger, A., Loutre, M.F., Galleé, H., 1998. Sensitivity of the LLN climate model to the astronomical and CO2 forcings over the last 200 ky, Climate Dynamics 14(9), 615–629. Berger, A., Li, X.S., Loutre, M.F., 1999. Modelling northern hemisphere ice volume over the last 3 Ma, Quaternary Science Reviews 18, 1–11. Berger, W.H., Bickert, T., Yasuda, M.K., Wefer, G., 1996. Reconstruction of atmospheric CO2 from ice-core data and the deep-sea record of Ontong Java plateau: the Milankovitch chron., Geol. Rundsch. 85, 466–495. Crucifix, M., Loutre, M.F., Lambeck, K., Berger, A., 2001. Effect of isostatic rebound on modelled ice volume variations during the last 200 kyr, Earth and Planetary Science Letters 184(3–4), 623–633. Crucifix M., Loutre M.F., Tulkens Ph., Fichefet T., Berger A., 2002. Climate evolution during the Holocene: A study with an Earth system model of intermediate complexity, Climate Dynamics 19, 43–60. Dansgaard, W., Johnsen, S.J., Clausen, H.B., DahlJensen, D., Gundestrup, N.S., Hammer, C.U., Hvldborg, C.S., Steffensen, J.P., Sveinbjo¨rnsdottir, A.E., Jouzel, J., Bend, G., 1993. Evidence for general instability of past climate from a 250kyr ice-core record, Nature 364, 218–220. EPICA Community Members, 2004. Eight glacial cycles from an Antarctic ice core. Nature 429 (6992), 623–628. Galleé, H., van Ypersele, J.-P., Fichefet, T., Tricot, C., Berger, A., 1991. Simulation of the last glacial cycle by a coupled, sectorially averaged climate-ice sheet model. Part I: The climate model, Journal

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of Geophysical Research-Atmospheres 96, 13 139–13 161. Galleé, H., van Ypersele, J.-P., Fichefet, T., Marsiat, I., Tricot, C., Berger, A., 1992. Simulation of the last glacial cycle by a coupled, sectorially averaged climate ice sheet model. Part II: Response to insolation and CO2 variation, Journal of Geophysical Research-Atmospheres 97, 15 713–15 740. Galleé, H., Berger, A., Shackleton, N.J., 1993. Simulation of the climate of the last 200 kyr with the LLN 2-D model, in: W.R. Peltier (Ed.), Ice in the Climate System, NATO ASI Series I, Global Environmental Change 12, 321–341, Springer, Berlin. Hays, J.D., Imbrie, J., Shackleton, N.J., 1976. Variations in the earth’s orbit: pacemaker of the ice ages, Science 194, 1121–1132. Imbrie, J., Hays, J., Martinson, D.G., McIntyre, A., Mix, A.C., Morley, J.J., Pisias, N.G., Prell, W.L., Shackleton, N.J., 1984. The orbital theory of Pleistocene climate: support for revised chronology of the Marine 18O record, in: A. Berger, J. Imbrie, J.D. Hays, G. Kukla, B. Saltzman (Eds.), Milankovitch and Climate, Reidel, Dordrecht, 269–305. Jouzel, J., Barkov, N.I., Barnola, J.M., Bender, M., Chappellaz, J., Genthon, C., Kotlyakov, V.M., Lorius, C., Petit, J.R., Raynaud, D., Raisbeck, G., Ritz, C., Sowers, T., Stievenard, M., Yiou, F., Yiou, P., 1993. Extending the Vostok ice-core record of paleoclimatic to the penultimate glacial period, Nature 364(6436), 407–412. Kubatzki, C., Claussen, M., 1998. Simulation of the global bio-geophysical interactions during the Last Glacial maximum, Climate Dynamics 14, 461–471. Li, X.S., Berger, A., Loutre, M.F., 1998. CO2 and northern hemisphere ice volume variations over the

middle and late quaternary, Climate Dynamics 14(7–8), 537–544. Loutre, M.F., Berger, A., 2003. Stage 11 as an analogue for the present interglacial, Global and Planetary Change 36, 209–217. Loutre, M.F., Berger, A., Dutrieux, A., Galleé, H., 1994. The response of the LLN climate model to the astronomical forcing over the last glacial– interglacial cycle, Terra Nostra, Schriften der Alfred-Wegener-Stiftung 1/94, 11–15. Martinson, D.G., Pisias, N.G., Hays, J.D., Imbrie, J., Moore, J.T.C., Shackleton, N.J., 1987. Age dating and the orbital theory of the ice ages: development of a high-resolution 0 to 300 000-year chronostratigraphy, Quaternary Research 27, 1–29. Parrenin, F., Re´my, F., Ritz, C., Siegert, M.J., Jouzel, J., 2004. New modeling of the Vostok ice flow line and implication for the glaciological chronology of the Vostok ice core. Journal of Geophysical Research 109, D20102, DOI: 10.1029/ 2004JD004561. Petit, J.R., Jouzel, J., Raynaud, D., Barkov, N.I., Barnola, J.M., Basile, I., Bender, M., Chappellaz, J., Davis, M., Delaygue, G., Delmotte, M., Kotlyakov, V.M., Legrand, M., Lipenkov, V.Y., Lorius, C., Pe´pin, L., Ritz, C., Saltzman, E., Stievenard, M., 1999. Climate and atmospheric history of the past 420 000 years from the Vostok ice core, Antarctica, Nature 399(6735), 429–436. Raynaud, D., Barnola, J.M., Souchez, R., Lorrain, R., Petit, J.R., Duval, P., Lipenkov, V.Y., 2005. Revisiting the Vostok record: the CO2 paradox of marine isotope stage 11. Shackleton, N.J., Le, J., Mix, A., Hall, M.A., 1992. Carbon isotope records from Pacific surface waters and atmospheric carbon dioxide. Quaternary Science Reviews 11(4), 387–400.

Section 2 Methods of Palaeoclimate Reconstruction and Dating (ed. Frank Sirocko)


5. Introduction – Palaeoclimate Reconstructions and Dating Frank Sirocko Institute for Geosciences, Johannes Gutenberg-University, Becherweg 21, 55099 Mainz, Germany

This book on the ‘Climate of Past Interglacials’ concentrates on the last four interglacials preceding the Holocene, that is, during the time from about 100 to 450 kyr. The first knowledge on this time has been extracted almost 40 years ago from marine deep sea cores with a sampling resolution of several thousand of years. Deepocean sediments can be rather easily dated, because 18 O stratigraphy (Shackleton and Opdyke, 1973) can be applied all over the ocean basins and tuned directly to the ice volume/sea-level master curve of SPECMAP (Martinson et al., 1987), which is based on the beat of the orbital insolation cyclicities (Berger, 1978; Berger et al., this volume). Sediments of the ocean are today investigated with a much higher time resolution (see for example the papers by Vautravers et al., Roucoux et al. and Rein et al. in this volume) and have revealed century-scale abrupt events of changing wind directions, sea-surface temperatures, strength of ocean currents, ice-sheet stability and deep-water formation, most of which are processes that can directly affect climate in the regions under their direct influence or even worldwide (e.g. El Ninõ or the thermohaline circulation). The stratigraphy of such rapid events is, however, derived from a correlation with ice core chronologies. In particular, the GRIP, GISP2 and NorthGRIP drillings in Greenland (Johnsen et al., 1992; Grootes et al., 1993; NorthGRIP, 2004) have shaped our knowledge about the very high speed of climate change on the northern hemisphere and serve now as master chronology for MIS 1–5. The long records of Vostok (Petit et al., 1999) and EPICA (EPICA Community Members, 2004) from Antarctica

are the most highest resolution source of information for the southern hemisphere during the last 800 000 years. The stratigraphy of these long cores is, however, at some tie points tuned to SPECMAP. Thus, the stratigraphy of ice cores and marine sediments is based on the robustness of the oxygen isotope stratigraphy which was first developed by Emiliani (1955) and established by Nicholas Shackleton (1937–2006), to whom this volume on the ‘Climate of Past Interglacials’ is dedicated. Latest research of this topic goes even further by modelling oxygen isotope variations. See also the compilation of several different approaches for sea-level reconstructions by Siddall (this volume). The reliability of the oxygen isotope stratigraphy was established by numerous radiometric dates, mainly 14C for the last 55 000 years, U/Th for the time back to 350 000 years and K/Ar or Ar/Ar datings well back into the early Pleistocene. All of these methods have specific demands on the amount of datable material and are prone to analytical error. The time of interest for this book is from 100 to 450 kyr, thus mainly in the range of U/Th dating, and we have thus two papers in this chapter that deal with this isotope system. Scholz et al. (this volume) developed a refinement of coral ages for a sea-level highstand during isotopic stage 7a, and Frechen (this volume) applied the U/Th technique to terrestrial peats in North Germany, where the age of the Holsteinian (MIS 9 or 11) has come under strong debate (see also the paper by Geyh and Mu¨ller, this volume). Both of these applications depend on the immobility of U and Th in peat or coral aragonite.

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Whether these prerequisites for reliable dating are indeed given is a serious matter of concern; see for example the U/Th values for the classical Eemian profile of Gro¨bern. The U/Th values presented by Frechen et al. (this volume) would date the Gro¨bern site into MIS 7e. This is highly unlikely, and Frechen et al. themselves claim that Gro¨bern should be MIS 5e. Accordingly, U/Th dates on peat have to be interpreted with some suspicion, which has to be kept in mind before the U/Th dating of the Holsteinian into MIS 9 can be indeed taken for granted (Geyh and Mu¨ller, this volume). The high demand for dating techniques for terrestrial sediments of the last 400 000 years led Krbetschek and Degering (this volume) to develop the new luminescence dating technique ‘radioluminescene’, which can be applied to clastic feldspar minerals from well-bleached sedimentary environments. They apply this technique to the sediments of Munster (North Germany) and arrive at an age of 334 kyr for the top of the Kieselgur, again indicating an age of MIS 9 for the Holsteinian. The dating of the diatomite at Munster is particularly important, because this is the only clearly Holsteinian record which is varved and has been used to determine the length of this interglacial to about 15 000 years (Mu¨ller, 1974). It is not only of academic interest whether the Holsteinian belongs to MIS 9 or MIS 11, because MIS 11 is the only interval during the late Pleistocene when the insolation forcing was similar to that during the Holocene. MIS 11 was also the longest of the last five interglacials (Fig. 5.1, Petit et al., 1999, see also McManus et al., 2002). The Munster record shows two spikes of Betula and Pinus dominance during the interglacial, which was interpreted as evidence for the occurrence of extreme cold events during an interglacial (Mu¨ller, 1974). Thus, it is of specific importance whether Munster indeed represents MIS 11, because it would be the only annually varved record for an older interglacial with climate forcing comparable to the Holocene and thus could inform us

about the processes and rates of climate change, which may lie ahead of us at the end of the Holocene. The most established continental record of the late Pleistocene interglacials is the VOSTOK ice core (Petit et al., 1999). We use these data for Fig. 5.1, even if the absolute dates of the interglacials have slightly changed in the new EPICA chronology. Based on the Vostok chronology, all the five interglacials have apparent differences in length, with MIS 11 of more than 20 000 years being the longest, and MIS 7e with less than 5000 years being the shortest. MIS 1, 5e and 9e are all about 12 000 years long. MIS 7e is spike shaped, MIS 5e and 9e are sawtooth shaped, quite in contrast to MIS 1 and MIS 11 which show no asymmetric shape. Another feature of importance for the forcing of past interglacials is documented in the trace gas content of the Vostok ice core (Fig. 5.1). Methane parallels the D (temperature) trend almost perfect. CO2 in contrast lags the temperature record by 4000 years during the last glacial inception (MIS 5e–5d transition), and a lag is also clearly visible for the glacial inception at the end of MIS 11 and MIS 9e. Accordingly, some parts of the global interglacial climate systems remained on an interglacial level much longer than the temperature over Antarctica. This pattern was the same during all past interglacials except MIS 7e. We do not want to speculate on possible reasons for this observation, but it matches the observation of several authors in this book, that the floral changes on the northern hemisphere developed time transgressive during the last interglacial. The CO2 lag as well as its relation to CH4 changes is a particularly important point, because it hints at interaction processes between various climate system components (see Claussen, this volume), in particular between the terrestrial biosphere, wetlands, marine biosphere and oceanic chemistry, which are yet not understood. Presumably, CO2 changes amplify climate changes – a process which becomes interesting in the light of the present-day climate

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changes in which anthropogenic CO2 emissions have to be regarded as forcing of the climate system. The CO2 lag has, however, also implications for the dating and duration of past interglacials. The duration of the interglacial CO2 maximum above the value of 118 kyr, which marks the inception of the last glacial (see Seelos and Sirocko, this volume; Kubatzki et al., this volume), is 18 300 years, the length of the interglacial temperature maximum only 14 800 years (Fig. 5.1). Apparently, the duration of an interglacial depends on a definition, which is classically based on the oxygen isotope stratigraphy and maximum sea-level highstands (see above). Martinson et al. (1987) used the maximum gradients of sea-level increase or decrease to define the boundaries of the SPECMAP stages. This approach is still used for all marine sediment cores. Terrestrial palynologists, however,

prefer the expression of a ‘thermomere’ to address a phase with abundant thermophilous pollen, and this can be an interglacial or an interstadial. The ice core community, in contrast, has recently presented a new approach to define an interglacial, i.e. they used the lowest Holocene D values in the EPICA core (EPICA Community Members, 2004) and address all past intervals with D values above this Holocene minimum value as ‘interglacial’. For Fig. 5.1, we used the midHolocence CO2 concentration of 260 ppm to mark a typical interglacial value, although peak values of 280 ppmv or 300 ppmv might be the maximum natural values during the Holocene. The above examples show that the name ‘interglacial’ is indeed used in very different ways. The CO2 lag (see above) indicates clearly that different parts of the climate system stay on interglacial values for

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different amounts of time, and these differences can last up to several thousand years (Fig. 5.1). The controversial discussion about the length of past interglacials (Kukla et al., 1997) versus (Turner, 2002) may thus indeed be obsolete, because an interglacial has apparently different durations at different locations even on the same hemisphere. The principal asynchroneity (see-saw) of the MIS 3 interstadials between the northern and southern hemisphere was demonstrated very convincingly by Blunier et al. (1998). Asynchroneities between the low latitudes and the northern hemisphere have been reported from an early sea-level rise (Henderson and Slowey, 2000) and early highstands of African lakes (Trauth et al., 2003) at the beginning of Termination II. Steep gradients in the vegetation belts across Europe during the last glacial inception are just about to be detected. Sańchez Gon˜i et al. (2005) and Sirocko et al. (2005) demonstrated on the basis of highresolution records from the Portugal margin and Eifel maar lakes in soutwest Germany respectively that the development of vegetation at the end of the Eemian in Portugal/ France/southwest Germany is out of phase with the floral evolution in north/east Germany and Scandinavia. Both authors claim for steep gradients in the SST fields of the Atlantic Ocean, which could decouple the climate of northern Europe from the climate in southern Europe. Mu¨ller and Sańchez Gon˜i (this volume) discuss the vegetation evolution of southern Germany in the same context. Seelos and Sirocko (this volume) develop a correlation between cold events in the North Atlantic; a north German dust record, and dust records from Eifel maar sediments. Again, the floral evolution between north Germany and the Eifel is offset by several thousand years when records are correlated on the basis of dust events. The C-events are probably the best indicator for the existence of the North American ice sheet during the early Weichselian glaciation, but they can only occur after the ice has been already build up; thus they

cannot be used to date the last glacial inception at the end of the Eemian. The most direct indicator for the beginning of the last glaciation at 118 kyr is probably the beginning of sea-level regression as documented by U/Th dating of reef terraces (Lambeck and Chappell, 2001). The same age of 118 kyr is given by the end of the late Eemian speleothem growth in the Spannagel cave of the Alps (Holzka¨mper et al., 2004; Spo¨tl et al., this volume). This cave presents probably the best record for absolute dating of past interglacials in general, because it lies at the altitude of the modern snow line, which implies that the dripping water necessary for the formation of the speleothems freezes as soon as the snow line (representing an annual average temperature of 0 C) drops below the modern value. Time transgressive climate evolution is also documented by Rein et al. (this volume) for the ENSO region off Peru. Low-latitude climate shifts at the beginning and end of the last interglacial most probably lead the climate of the high northern latitudes, because the ENSO system is more directly forced by the seasonal gradient between spring and autumn climate (Clement et al., 1999) and not directly linked to the summer insolation forcing of the northern hemisphere. Such lead/lag relations are often used to depict causal mechanisms in the climate system, but it must not necessarily imply a forcing of high latitudes by lowlatitude processes (Claussen et al., 2003). The above evaluation follows the records presented in this book and is not a summary of the global available information on timing and forcing of past interglacials. There are many other important and excellent records worldwide, and a full review of the state of art on past interglacial research is beyond the scope of this book. The records presented are, however, completely sufficient to draw a few general inferences: 1. The beginning, end and duration of the past interglacials were not synchronous all over the world, i.e. parts of the climate

Palaeoclimate Reconstructions and Dating

system have been longer in an interglacial state than others. 2. Beginning and end of interglacials in the low latitudes and in the Antarctic lead respective changes on the northern hemisphere. 3. Time transgressive climate shifts are also strong over Europe, where the SST patterns of the North Atlantic drift presumably cause a stepwise shift of the vegetation zones at least during the end of the past interglacial, with a longer interglacial in southern Europe and shorter interglacial in the north. This dynamic climate evolution of the past interglacial must have been of high importance also for the evolution of mankind. Neanderthal hominids lived and hunted in Europe during the Eemian (Wenzel, this volume), but were replaced during MIS 3 by the modern humans. Genetic data indicate that these early humans migrated out of Africa during MIS 5 and lived in the Mediterranean during MIS 4 and the early MIS 3. Why and when they exactly moved into Northern Europe and Asia has certainly no relation to the past interglacial climate. The reason why they left Africa during the early MIS 5, however, might well have to been seen in this context. REFERENCES Berger, A. (1978). Long-term variations of daily insolation and Quaternary climatic changes. Journal of Atmospheric Sciences 35(12), 2362–2367. Blunier, T., Chappellaz, J., Schwander, J., Da¨llenbach, A., Stauffer, B., Stocker, T.F., Raynaud, D., Jouzel, J., Clausen, H.B., Hammer, C.U., and Johnsen, S.J. (1998). Asynchrony of Antarctic and Greenland climate change during the last glacial period. Nature 394, 739–743. Claussen, M., Ganopolski, A., Brovkin, V., Gerstengarbe, F.-W., and Werner, P. (2003). Simulated global-scale response of the climate system to Dansgaard-Oeschger and Heinrich events. Climate Dynamics 21, 361–370. Clement, A.C., Seager, R., and Cane, M.A. (1999). Orbital controls on the El Nino/southern

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oscillation and tropical Pacific climate during the last millennium. Nature 424, 271–276. Emiliani, C. (1955). Pleistocene temperatures. Journal of Geology 63, 538–578. EPICA Community Members (2004). Eight glacial cycles from an Antarctic ice core. Nature 429, 623–628. Grootes, P.M., Stuiver, M., White, J.W.C., Johnsen, S., and Jouzel, J. (1993). Comparison of oxygen isotope records from the GISP2 and GRIP Greenland ice cores. Nature 366, 552–554. Henderson, G.M., and Slowey, N. (2000). Evidence from U–Th dating against Northern Hemisphere forcing of the penultimate deglaciation. Nature 404, 61–66. Holzka¨mper, S., Mangini, A., Spo¨tl, C., and Mudelsee, M. (2004). Timing and progression of the last interglacial derived from a high alpine stalagmite. Geophysical Research Letters 31, L07201, doi:10.1029/2003GL019112. Johnsen, S.J., Clausen, H.B., Dansgaard, W., Fuhrer, K., Gundestrup, N., Hammer, C.U., Iversen, P., Jouzel, J., Staufer, B., and Steffensen, J.P. (1992). Irregular glacial interstadials recorded in a new Greenland ice core. Nature 359, 311–313. Kukla, G., McManus, J.F., Rousseau, D.-D., and Chuine, I. (1997). How long and how stable was the last interglacial? Quaternary Science Reviews 16, 605–612. Lambeck, K., and Chappell, J. (2001). Sealevel change through the last glacial cycle. Science 292, 679–685. Martinson, D.G., Pisias, N.G., Hays, J.D., Imbrie, J., Moore Jr., T.C., and Shackelton, N.J. (1987). Age dating and the orbital theory of the ice ages: Development of a high-resolution 0 to 300,000-year chronostratigraphy. Quaternary Research 27, 1–29. McManus, J.F., Oppo, D.W., Keigwin, L.D., Cullen, J.L., and Bond, G.C. (2002). Thermohaline circulation and prolonged interglacial warmth in the North Atlantic. Quaternary Research 58, 17–21. Mu¨ller, H. (1974). Pollenanalytische Untersuchungen mit Jahresschichtza¨hlungen an der holststeinzeitlichen Kieselgur von Munster Brehloh. Geologisches Jahrbuch A 21, 107–140. NorthGRIP. (2004). High resolution record of Northern Hemisphere climate extending into the last interglacial period. Nature 43, 147–151. Petit, J.R., Jouzel, J., Raynaud, D., Barkov, N.I., Barnola, J.M., Basile, I., Bender, M., Chappellaz, J., Davis, J., Delaygue, G., Delmotte, M., Kotlyakov, V.M., Legrand, M., Lipenkov, V.M., Lorius, C., Pe`pin, L., Ritz, C., Saltzman, E., and Stievenard, M. (1999). Climate and atmospheric history of the past 420,000 years from the Vostok ice core, Antarctica. Nature 399, 429–436. Raynaud, D., Barnola, J.-M., Souchez, R., Lorrain, R., Petit, J.-R., Duval, P., and Lipenkov, V. (2005). The record of marine isotopic stage 11. Nature 436, 39–40.

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Sańchez Gon˜i, M.F., Loutre, M.F., Peyron, O., Santos, L., Duprat, J., Malaize, B., Turon, J.-L., and Peypouquet, J.-P. (2005). Increasing vegetation and climate gradient in Western Europe over the last glacial inception (122–110 ka): data model comparison. Earth Planetary Science Letters 231, 111–130. Shackleton, N.J., and Opdyke, N. (1973). Oxygene isotope and paleomagnetic stratigraphy of equatorial Pacific core V28-238: Oxygene isotope temperatures and ice volume on a 105 and 106 * 10 year scale. Quaternary Research 3, 39–55.

Sirocko, F., Seelos, K., Schaber, K., Rein, B., Dreher, F., Diehl, M., Lehne, R., Ja¨ger, K., Krbetschek, M., and Degering, D. (2005). A Late Eemian Aridity Pulse in central Europe during the last glacial inception. Nature 436, 833–836. Trauth, M.H., Deino, A.L., G.N., B., and Strecker, M.R. (2003). East African climate change and orbital forcing during the last 175 kyr BP. Earth and Planetary Science Letters 206, 297–313. Turner, C. (2002). Problems of the duration of the Eemian interglacial in Europe North of the Alps. Quaternary Research 58, 45–48.

6. Late Quaternary Interglacials in East Antarctica From Ice-Core Dust Records Barbara Delmonte1,2, Jean Robert Petit1, Isabelle Basile-Doelsch3, Emil Jagoutz4 and Valter Maggi2 1

LGGE-CNRS, BP96, 38402, Saint Martin d’He`res, France University of Milano-Bicocca, DISAT, Piazza della Scienza 1, 20126 Milano, Italy 3 CNRS-CEREGE, UMR 6635, Europole Me´diterraneén de l’Arbois, BP 80, 13545, Aix en Provence, France 4 Max Planck Institute of Chemistry, Kosmochemistry Department 55020 Mainz, Germany 2

ABSTRACT Aeolian dust records from deep East Antarctic ice cores evidence extremely low dust fluxes during the last five interglacials (10 to 25 times lower than in glacial periods), related to reduced primary production and mobilization on the Southern Hemisphere continents, to changes in atmospheric transport and hydrological cycle. The Sr–Nd isotope fingerprint of aeolian dust in Antarctica suggests a dominant southern South America provenance during Quaternary glacial times, but the first geochemical data for Stage 5.5 and the Holocene presented in this work show significant differences and open the possibility for a different source mixing. Dust-size variability in the EPICA-Dome C ice core suggests shorter transport time for dust or more direct air mass penetration to the site during interglacials with respect to cold periods and a clear multisecular scale mode of atmospheric circulation variability during the Holocene. 6.1 INTRODUCTION Mineral aerosol (dust) deflated from continental areas and transported in the atmosphere is of importance for the climate system (e.g. Harrison et al., 2001; Houghton et al., 2001 and references therein) for biogeochemical cycles and can be used as tracer

for depicting atmospheric circulation patterns and variability. Dust affects the solar radiation. The radiative impact of small (< 20 mm in diameter) dust particles is nonlinearly related to a number of factors such as their atmospheric concentration and vertical distribution, optical and aerodynamic properties as well as mineralogical composition (e.g. Sokolik and Toon, 1996; Claquin et al., 1998). Dust plays a role in atmospheric chemistry (see Harrison et al., 2001 for a general overview) and processes as formation of cloud droplets (Zhang and Carmichael, 1999) and ice nuclei (e.g. Rogers and Yau, 1989), as well as in biogeochemical cycles through nutrients supply to terrestrial (e.g. Swap et al., 1992) and marine ecosystems (e.g. Falkowski et al., 1998; Hutchins and Brunland, 1998; Fung et al., 2000), potentially influencing the global carbon cycle and the atmospheric concentration of greenhouse gases. Soil dust emissions at present time are estimated around 2150 Mt=yr on global scale (Houghton et al., 2001), with high spatial and temporal variability. In particular, emissions in the Southern Hemisphere, which is devoid of major dust sources, are less than one-fifth of those from the Northern Hemisphere, where North Africa, the Middle East, central Asia and the Indian subcontinent play a major role (Prospero et al., 2002). The mineral particles deflated from the austral continental landmasses are of

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interest for air mass tracking as they can be transported long range into the mid-to-high troposphere through the zonal westerly wind flux, ultimately reaching the interior of Antarctica, where they are archived in the ice layers. The reconstruction of Quaternary ‘background’ dust (of nonvolcanic origin) variability and provenance from deep ice cores recovered from low accumulation sites of the high East Antarctic Plateau can therefore provide detailed information about palaeoenvironmental conditions at the dust-source regions, air mass exchanges and variability between mid and high latitudes of the southern hemisphere and past atmospheric circulation changes on timescales from glacial–interglacial cycles (e.g. Petit et al., 1999) to submillennial, or centennial periods (Delmonte et al., 2004b, 2005). Amongst ice cores, the Vostok ice core (Fig. 6.1) (78 S, 106 E, 3480 m a.s.l.) first provided climate records spanning the last 420 kyr (Petit et al., 1999), followed by the 340 kyr long Dome Fuji record (Watanabe et al., 2003). Today, the EPICA (European Project for Ice Coring in Antarctica) deep ice core from Dome C (East Antarctica, 75 069S, 123 219E, 3233 m a.s.l.) allowed the extension of the climate record back to 740 kyr BP, corresponding to the last eight glacial cycles (EPICA Community Members, 2004). At shorter timescale, the investigation of the timing and magnitude of Holocene climatic variability has today become a major focus for the comprehension of present-day climatic trends and the assessment of anthropogenic contributions on natural climate change (Houghton et al., 2001). Over the last decade, a number of marine sediment cores from the North Atlantic (e.g. Bond et al., 1997; Bianchi and McCave, 1999), from the Indian Ocean (e.g. Sakar et al., 2000) and the South-East Pacific (e.g. Lamy et al., 2001), as well as atmospheric proxies from tropical latitudes (Moy et al., 2002) and northern polar areas (Mayewski et al., 1997) highlighted a pronounced millennial scale mode of variability of atmospheric and oceanic indicators over the

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Holocene epoch. High temporal resolution analyses followed and indicated that the millennial scale structures from North Atlantic deep-sea sediment records are actually composed by high-frequency oscillations (Bond et al., 2001). An 1500-year oscillation of climate was suggested and possibly associated with solar activity. At the high latitudes of the Southern Hemisphere, secular scale periodicities (of 200and 400-year duration) were detected (Leventer et al., 1996; Domack and Mayewski, 1999; Domack et al., 2001) in Holocene biogeochemical sediment records from the Antarctic Peninsula. The changes in sediment properties were interpreted in terms of local changes in upper ocean conditions. Aeolian dust preserved in Antarctic ice cores has the potential to document climate variability at different timescale; also, it can provide relevant information that cannot be inferred from other proxies. Indeed, for the

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mineral aerosol spreads over Antarctica, the narrow and geographically fixed location of their main source regions (southern South America see §3), their relatively longdistance transport and their low chemical reactivity together make this proxy an almost passive tracer worth capturing the changes in atmospheric circulation patterns around Antarctica. Moreover, the size distribution of the dust is sensitive to the transport conditions, and it could provide an additional information on the pathway. This property is likely unique by comparison to other climatic indicators (e.g. stable isotope composition of ice or total concentration of chemical component) having extensive and variable sources (e.g. austral ocean) or more sensitive or reactive to other factors (e.g. temperature, vapour saturation pressure and other chemicals) during their pathway. In the following sections, we first present the broad patterns of aeolian dust variability in East Antarctica during the late Quaternary (over the last four glacial cycles and last 0.4 million years) within the global context of glacial/interglacial cycles (§2) and with particular focus on the characteristics of aeolian dust during interglacials. Then, we summarize the present-day knowledge about geographic provenance of dust to the Antarctic snow during glacial times and provide evidence for a possible source differentiation between cold glacials and warm interglacials (section 6.3). Afterwards, we present some recent findings on the spatiotemporal variability of dust transport in East Antarctica for two periods of interest: the last glacial maximum (LGM) to Holocene climatic transition (section 6.4) and the Holocene (section 6.5).

6.2 LATE QUATERNARY AEOLIAN DUST VARIABILITY IN ANTARCTICA 6.2.1 Glacial/interglacial cycles The late Quaternary stable isotope record from EPICA-Dome C (hereinafter EDC) ice

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core (EPICA Community Members, 2004, Fig. 6.2a), taken as proxy for temperature variations, shows a sawtooth pattern of warm interglacial stages (and corresponding to marine isotopic stages – MIS – 11.3, 9.3, 7.5, 5.5 and to the Holocene) followed by glacial periods increasingly colder and punctuated by cool interstadials. The EDC isotope profile is very similar to the previous Vostok (Petit et al., 1999) and Dome Fuji (Watanabe et al., 2003) records, demonstrating a rather good uniformity of glacial– interglacial climate changes across Antarctica. Temperature changes highlighted also the correlation with atmospheric greenhouse gases (CO2 and CH4) content, the most rapid changes occurring during glacial–interglacial transitions (Petit et al., 1999; EPICA Community Members, 2004). The pattern of dust concentration variability from EDC (Fig. 6.2b) and Vostok (Fig. 6.2c) ice cores measured by Coulter Counter technique on discrete ice samples is very similar and is anti-correlated at first order with the isotope record. Interglacial periods are characterized by extremely low dust inputs ð0:40:6 mg m2 yr1 Þ, while 10 to 25 times higher fluxes characterize cold glacial stages (up to 50 times and even higher for concentrations, due to the reduced precipitation rate). This evidence points out that the major mode of glacial/interglacial aeolian dust variations in central Antarctica is remarkably uniform. Indeed, such pattern is of global significance. There is considerable evidence for enhanced atmospheric dust load during Quaternary cold periods, and consequently higher deposition in oceans and on continents, the evidence coming from northern polar records (e.g. Steffensen et al., 1997; Ruth et al., 2003), from terrestrial (e.g. Kukla et al., 1990) and from marine (e.g. Rea, 1994) sequences. A few regional exceptions however do exist (see Kohfeld and Harrison, 2001 for an update overview), especially from tropical and subtropical latitudes, generally associated with opposite climate changes or with the variability of atmospheric pathways. For

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example, a 300-kyr long marine record from southwestern Africa, off the Namib Desert (20–25 S), highlighted humid

glacials and dry interglacials occurring in response to equatorward shifts of precipitation belts during cold stages (Stuut et al., 2002). Climate and atmospheric variability over the Quaternary is mainly related to the periodicity of the Earth’s orbital parameters (e.g. Imbrie et al., 1993) modulating the amount, as well as the seasonal and the latitudinal distribution of solar radiation (Berger and Loutre, 1991) and likely initiating climate changes. As for most climate proxies recorded in Antarctic ice cores (Petit et al., 1999 and EPICA Community Members, 2004), a large part of the dust and isotope variance is concentrated in the 100- and 41-kyr spectral bands, while their links with similar periodicities of the eccentricity and the obliquity respectively still need to be assessed. The overall higher global atmospheric dust load characterizing cold climatic stages can be explained by a number of synergetic factors: like (1) the widespread continental aridity and the lower atmospheric vapour content imposed by cooler temperatures, leading to changes in soil moisture (hence resistance to erosion) and vegetation cover; (2) the enhanced primary production of dust through physical weathering processes in periglacial areas (e.g. frost and thawing cycles), which became particularly important during cold periods because of the glacier cover extension in mountain area and large ice sheet cover on continental landmasses; (3) the substantial enlargement of the dust-source areas due to sea-level lowering; (4) the reduced intensity of the hydrological cycle, leading to less efficient scavenging processes by precipitations and consequently to an increase in the atmospheric dust residence time (Yung et al., 1996); and finally (5) a generally more vigorous atmospheric circulation associated with steeper latitudinal thermal gradients (e.g. COHMAP Project Members, 1988; Kohfeld and Harrison, 2001). Under the generally windier conditions prevailing in the continents in glacial times, processes


like soil deflation, dust injection into the atmosphere and long-range transport were very efficient. Under mild interglacial climates, in contrast, the dust production, mobilization and transport decreased on average because of damper environmental conditions, dense vegetation cover, enhanced evaporation, precipitation and aerosol scavenging. Dust and isotope records from Greenland and Antarctic ice cores differ significantly. Large and very rapid dust changes occurred in Greenland during the Dansgaard– Oeschger (D/O) events of the last glacial period, and during the last climatic transition, where the Younger Dryas event was accompanied by a rapid return of dust concentration to glacial levels (e.g. Ruth et al., 2003). Aeolian dust archived in Greenland mostly originates from central Asia (Biscaye et al., 1997), travels for several thousands of kilometres through the high troposphere over the Pacific Ocean and finally reach northern latitudes. As suggested by the dust/isotope correlation, the dust input to Greenland is strongly anti-correlated with local temperature. For East Antarctic records, the southern South America is the dominant source region of dust during cold periods (see section 6.3). Indeed, the climate of this region is sensitive to the sea ice cover, and it was greatly influenced by the extended sea ice cover in the South Atlantic Ocean during glacial times, modulating the intensity and average position of the polar front (e.g. Heusser, 1989). Altogether, this makes the East Antarctic dust records more sensitive to the source and to the southern South American climate and environment. Moreover, this peculiar area under the westerlies is close to the Drake Passage and the Weddell Sea, which are key ocean locations well connected to large-scale deep-ocean currents (e.g. North Atlantic Deep Water, circum polar current) and sensitive in turn to the global ocean circulation and climate. The apparent correlation between the global ice volume (Fig. 6.2c) and the dust records from East Antarctica

57

supports such a view, with possibilities for large-scale teleconnections. Interestingly, the Antarctic dust records display a glacial/interglacial pattern that is similar to the magnetic susceptibility record of loess–palaeosoil sequences (Kukla et al., 1990) from the Chinese Loess Plateau (Fig. 6.2d). Despite the records yet having to be synchronized, the similarity suggests a global character or at least an intercontinental (central Asia–southern South America) connection of climate and environment with respect to aeolian dust production, mobilization, transport and accumulation. 6.2.2 Characteristics of dust during interglacials Figure 6.3 displays the records from the EDC ice core for stable isotopes and aeolian dust during the last five interglacials (Holocene, MIS 5.5, MIS 7.5, MIS 9.3 and MIS 11.3). A sharp rise of the isotope content with a variable pattern marks the climatic transitions (Terminations) from glacial stages to interglacials of variable duration. If the time period spent by the isotope record above the 403 per mil level (300-year average minimum value observed during the full Holocene epoch) is considered as threshold for interglacial temperature conditions (EPICA Community Members, 2004), the duration of the last five interglacials (Fig. 6.3a and 6.3d) looks very variable around 16 9 kyr, with a maximum of 29 kyr (MIS 11.3) and a minimum of 5 kyr (MIS 7.5) During these warm periods, the dust concentration in the EDC ice core is always very low (Fig. 6.3b and 6.3e), typically lower than 30 ppb, corresponding to about 0:5 mg m2 yr1 in flux on average (Table 6.1). From one interglacial to the other, dust concentration and fluxes are very similar, and it seems that in all cases the low (interglacial) dust levels are reached a few thousand years before the culmination of the isotope to the interglacial value. Sometimes a two-step increase of isotope content occurs afterwards as in the case of the last climatic

58

Barbara Delmonte et al. Termination I 0 1

–360

80 –360

–400 2

4

Termination II

Termination III

Termination IV

Termination V

100 120 140 160 180

220 240 260 280 300

300 320 340 360 380

380 400 420 440 460

5.5

–360

–360

7.5

–360

–400

–400

–400

9.3

12

–440

–440

–440

(b)

1000

–1000

–1000

–1000

–1000

100

Holocene

–440

100

1

CPP (%)

(c)

25

25

20

20

15

15

10

10

20 40 60 80

25

7.5

25 8

20 6

100 120 140 160 180

30 ppb

10 11.3

9.3

25

20

20

15

15

15

10

10

10

220 240 260 280 300

12

100

10

2

0

5.5

10

100

10

10

10

10

8

100

6

–403‰

–400

–440

2

11.3

10

8

6

(ppb)

δD‰ (V-SMOW)

(a)

20 40 60

300 320 340 360 380

Coarse dust 12

Fine dust 380 400 420 440 460

Age (kyr BP)

Fig. 6.3 EDC records of stable isotope, dust concentration and dust-size changes during the last five climatic transitions. (a) – The deuterium record with indication (horizontal dashed line) of the minimum 300-year average value observed during the full Holocene epoch (403 per mil, EPICA Community Members, 2004). The D values above this level are taken as indicators for interglacial temperature conditions, and interglacials are highlighted by the yellow bands. (b) – EDC dust concentration changes (log scale). Typical interglacial levels lie below the 30 ppb level (dashed horizontal line). (c) – Dust size (expressed as ‘coarse particle percentage’, CPP, corresponding to the proportion of particles having a diameter between 3 and 5 mm with respect to the total dust mass, typically included in the interval of diameter between 1 and 5 mm). At the time of each climatic transition, the mean particle size progressively decreases (blue arrows) during glacials and until the very beginning of the climatic transition, then sharply increases (red arrows) at the onset of warm conditions. It is interesting to point out that Termination III displays two drastic changes of dust concentration accompanied by a minimum of dust size. One occurs around 260–270 kyr BP, during stage 8, the other around 250 kyr BP during the climatic transition.

transition and the Antarctic Cold Reversal (ACR) during Termination I (Fig. 6.3), a pattern possibly associated with the timing of final sea ice retreat in the South Atlantic (Jouzel et al., 1995). At the end of each glacial period, a systematic change in the particle grain size occurs at Dome C (Fig. 6.3c and 6.3f). The particle grade is given here by the coarse particle percentage, CPP parameter (see section 6.4). CPP always increases at the onset of warm periods and on average is higher during interglacials than during glacials (Table 6.1). During the last five terminations,

particle size displays a decreasing trend at the end of each glacial period, and in spite of some variability with spikes of large particles, it displays a minimum value at the very beginning of the climatic transition at all times (Fig. 6.3c). Such behaviour for EDC dust suggests that similar phenomena characterize the end of each cold period and the onset of each interglacial. This can ultimately provide information on atmospheric circulation patterns over Dome C. The occurrence of the same pattern during the last 400 kyr suggests this is a typical feature of climatic transitions in East Antarctica and

Late Quaternary Interglacials in East Antarctica 0

δD‰ (V-SMOW)

(d)

–360

10

Holocene

–380

120 –360

130 –360

5.5

–380

240

–380

250

320

–360

7.5

330

340

390 400 410 420 –360

9.3

–380

59

–400

–400

–400

–400

–420

–420

–420

–420

–420

–440

–440

–440

–440

100

100

80

80

Dust concentration (ppb, linear scale)

(e)

60

~16 kyr

100

40

40

20 0 30

60 16 ± 10

~29 kyr

100

80 22 ± 18

–403‰

–440 ~15 kyr

100

80

60

15 ± 8

~5 kyr

11.3

–380

–400

80 16 ± 7

60

60

40

40

40

20

20

20

20

0 30

0 30

0 30

0 30

25

25

25

25

25

20

20

20

20

20

15

15

15

15

15

10

10

10

10

10

19 ± 17

Coarse dust

CPP (%)

(f)

0

10

120

130

240

250

320

330

340

Fine dust 390 400 410 420

Age (kyr BP)

Fig. 6.3 (Continued) (d) – Isotope record (same as (a)) for interglacials. Horizontal blue arrows indicate the broad duration of each interglacial, calculated as the time period spent by the isotope record above the 403 per mil level. (e) – EDC dust concentration changes during interglacials reported in linear scale. For the Holocene and MIS 5.5, the high-resolution data from Delmonte et al. (2004a) are reported. For MIS 7.5, MIS 9.3 and MIS 11.3, the records are the same as reported in (b). Average concentration levels (red dot) are similar among interglacials. (f) – Dust-size (CPP) changes during interglacials. For the Holocene and MIS 5.5, the high-resolution data are reported; for older interglacials, a zoom of the profile (c) is reported.

likely reflects conditions of the ocean surrounding Antarctica, sea ice extent and its distribution as well as peculiar atmospheric circulation patterns. A closer inspection of mineral dust-size variability at Dome C (Fig. 6.3f), moreover, shows a marked mode of variability

throughout each interglacial. This can be better appreciated during MIS 5.5 and particularly during the Holocene (as it will be discussed in §5), which are documented at higher temporal resolution. As far as dust-size variability is concerned, it is very difficult to extrapolate

Table 6.1 Aeolian dust concentration and flux in East Antarctica during the last five interglacial stages from EDC ice core, with indication of dust-size changes Dust concentration (ppb) Dust flux ðmg=m2 yr1 Þ Glacial–interglacial dust size change

Holocene

MIS 5.5

MIS 7.5

MIS 9.3

MIS 11.3

15 8

16 10

22 18

16 7

19 17

0.4 (0.2–0.6)

0.4 (0.2–0.7)

0.6 (0.2–1)

0.4 (0.2–0.6)

0.5 (0.1–0.9)

Increase CPP (ca. þ10%)





The average dust levels reported in the table are calculated over the period when the isotope curve lies above the minimum 300-year average value observed during the Holocene epoch (403 per mil, EPICA Community Members, 2004, see Fig. 6.3). Taking such level as indicator for interglacial temperature conditions, the duration of warm periods looks very variable around 16 9 kyr on average between MIS 5.5 ð 16 kyrÞ, MIS 7.5 ð 5 kyrÞ, MIS 9.3 ð 15 kyrÞ and MIS 11.3 ð 29 kyrÞ.

60


results from EDC to the whole Antarctic since regional effects may play a fundamental role both during deglaciations (section 6.4.2) and during interglacials (section 6.5.2). Climate and dust records from other sites from the Plateau will be very instructive in this respect. 6.3 GEOGRAPHIC DUST PROVENANCE DURING GLACIALS AND INTERGLACIALS The source regions providing the bigger dust fluxes today are characterized, in general terms, by little or no ground cover, easily wind-erodible soils and seasonal wetness (Mahowald et al., 1999). A recent satellitebased worldwide geographical mapping of major atmospheric dust sources provided by Prospero et al. (2002) highlighted arid areas centred over topographical lows or on lands adjacent to topographical relief to constitute the major sources for dust transported long range. Indeed, terrains where the recent geomorphologic history has favoured the concentration of fine-grained material on low roughness surfaces are much more active suppliers than coarse-grained old sandy deserts already impoverished in the fine fraction (Tegen et al., 2002). Atmospheric transport exerts a strong size and mineralogical selection on particles, making small quartz

grains and plate clays, having the best aerodynamic properties, the most diffuse aeolian minerals around the globe (Gaudichet et al., 1992). The mineral dust deposited in ice on the East Antarctic Plateau after long-range transport is very fine grained (< 5 mm in diameter) and consists mostly of clays (mostly illite), crystalline silica, feldspars and minor amounts of pyroxenes and amphiboles, metallic oxides and volcanic glasses (Gaudichet et al., 1988, 1992). In order to depict the geographical provenance for aeolian dust at polar latitudes, Grousset et al. (1992) first proposed a successful geochemical approach, already used in oceanography (e.g. Grousset et al., 1988, Revel et al., 1996), consisting in the comparison of the 87Sr/86Sr versus 143Nd/144Nd isotopic signature of mineral particles extracted from Antarctic ice cores to that from potential source area (PSA) samples from the Southern Hemisphere (Table 6.2). The authors suggested a possible southern South American provenance for dust in East Antarctica during glacial stage 2 (MIS 2). Investigations on Antarctic dust provenance were largely developed later by Basile et al. (1997) and by Delmonte et al. (2004a, 2004b) on four East Antarctic sites from the plateau (Dome C, Vostok, Dome B and Komsomolskaya, Fig. 6.1). In parallel, the authors analysed a number of size-selected samples of loesses, aeolian deposits, sands

Table 6.2 87Sr/86Sr ð2 106 Þ, 143Nd/144Nd ð2 106 Þ and "Nd ð0Þ isotopic values for interglacial (Holocene and stage 5.5) samples from EDC and Vostok ice cores (1020 mg total dust per sample) Sample Holocene 1 – EDC Holocene 2 – EDC Holocene 3 – VK Holocene 4 – VK Stage 5.5 – EDC Stage 5.5 – VK Stage 5.5 – Volcanic

87

Sr/86Sr ð2 106 Þ 0.710013 (55) 0.709435 (37) 0.711200 (35) 0.709289 (50) 0.710213 (26) N.M. 0.704983 (36)

143

Nd/144Nd ð2 106 Þ 0.512407 (101) 0.512347 (95) 0.512126 (17) 0.512379 (44) 0.512211 (172) 0.512261 (75) 0.512823 (53)

"Nd ð0Þ 4:51 5:68 9:99 5:05 8:29 7:35 þ3:61

The whole analytical procedure is reported in Delmonte (2003) and in Delmonte et al. (2004a). Neodymium ratios are calculated as follows: "Nd ð0Þ ¼ ðð143 Nd=144 NdÞmeas =ð143 Nd=144 NdÞCHUR 1Þ 104 using the present-day CHUR (chondritic uniform reservoir) value for 143 Nd/144Nd ratio of 0.512638 (Jacobsen and Wasserburg, 1980). The measured 143Nd/144Nd were corrected for mass fractionation by normalizing to 146 Nd=144 Nd ¼ 0:7219, while the 87Sr/86Sr ratios were normalized to 86 Sr=88 Sr ¼ 0:1194.


and fluvioglacial sediments from key areas of the Southern Hemisphere and compared these to the ice core dust signature taking into account (1) size-dependent isotopic fractionation effects and (2) a possible 87 Sr=86 Sr shift for carbonate contribution (Delmonte et al., 2004a). The results of analysis are shown on Fig. 6.4. It can be observed that the isotopic signature for southern South America, constructed by collecting samples from the Argentine Pampas, Patagonia and other areas southward of 32 S, nicely fits that of ice core dust during glacial stages, suggesting southern South America was likely the dominant dust source during cold climatic stages for all the East Antarctic sites investigated. In particular, the South America samples from the Pampas region as well as those from the Cordoba loess region and from Patagonia south of 41 S show the best fit with glacial dust. The Sr–Nd isotopic fields for New Zealand and for the Antarctic Dry Valleys, however, also show a partial overlap with the South American and glacial dust ones (Fig. 6.4a), but complementary arguments allowed estimating their possible contribution as being negligible (Delmonte et al., 2004a). As far as New Zealand is concerned, the absence of tephra layers from the active Taupo Volcanic Zone in the Vostok ice core for the last 420 000 years (Basile et al., 2001), the very limited surface of both the North and South Island and the absence of significant contributions from the nearby large deserts of Australia suggest this region was unlikely to be the major dust source during cold periods. The Dry Valleys of coastal Antarctica were also unlikely the principal dust sources in glacial times: icefree areas were less extended than today during cold periods and the glacier’s fronts were closer to the coast (Denton and Hughes, 2002). Moreover, the primary production of fine-grained mineral particles was low in that region and limited to the mechanical alteration of rocks as the seasonal temperature variations and the hydrological cycle were reduced. Also, salts

61

deriving from the first stages of chemical weathering of rocks, like gypsum and carbonates, are very common in the Antarctic Dry Valleys (Campbell and Claridge, 1987), but they were not found in Antarctic ice (De Angelis et al., 1992). Finally, the strong catabatic winds blowing off the East Antarctic Plateau would lead to carry the mineral aerosol towards the ocean (Schwerdtfeger, 1984), and therefore a transport from the coast to the high Plateau would imply uplift of dust into the middle-high troposphere and a backward transport into the Continent, which seems quite unlikely. Using the same laboratory procedure as in Delmonte et al. (2004a), the Sr–Nd isotopic composition of EDC and Vostok dust from interglacial ice sections was measured and is reported in this work (Fig. 6.4a, 6.4b and Table 6.2). As the total amount of dust extracted from interglacial samples was extremely low (about 10 to 20 mg per sample), the error of measurement for Nd is relatively important. Despite this, and the limited number of samples analysed, the isotopic signature for interglacial dust looks significantly less radiogenic in 143Nd/144Nd ð10 < "Nd ð0Þ < 4:5Þ than glacial dust (Fig. 6.4b), with the exception of one sample from stage 5.5 containing volcanic ash, which shows obviously the typical fingerprint of volcanic rocks. The ice core dust isotopic data appear nicely aligned along a mixing line between two hypothetical volcanic and crustal end-members. The isotopic difference between samples from glacial and interglacial periods could be hardly accounted for by the South America dust collected south of 32 S and is reported in Fig. 6.4a. Indeed, environmental changes occurring within a given region and related to chemical weathering processes and/or pedogenesis would likely affect the Rb–Sr system (which is open and allows different elemental fractionation of Sr and Rb during chemical weathering leading to 87Sr/86Sr enrichment in the clay fraction) rather than the Sm–Nd system (Basile et al., 1997; Delmonte, 2003).

62


II

VK glacial stages 4 to 12 EDC glacial stages 2, 4 and 6 DB glacial stage 2 KMS Termination I Old Dome C stage 2 EDC and VK interglacials [Holocene, Stage 5.5] EDC volcanic dust sample

I

South America

ε

8

Sr = 0

12

4

East Antarctic aeolian dust (glacials) 9

0

GLACIAL DUST New Zealand (South Island)

–12

Adelie land

Dry Valleys

0

NVL (Antarctica)

Af

–3

–9

Hol-EDC

Hol-VK Hol-EDC

–6

a ric

East Antarctic aeolian dust (interglacials)

–20

3

(Antarctica)>>>

h ut So

–16

Australia

εNd (0)

–8

New Zealand (North Island)

εNd (0)

–4

6

5.5-VK

INTERGLACIAL DUST

5.5-EDC Hol-VK

–24

IV

III 0.700

(a)

0.710

0.720

0.730

0.740

–12

0.750

87Sr/86Sr

0.704

(b)

0.706

0.708

0.710

0.712

0.714

87Sr/86Sr

Fig. 6.4a 87Sr/ 86Sr versus "Nd ð0Þ isotopic signature of East Antarctic ice core dust from glacials (white open circles, from Grousset et al., 1992; Basile et al., 1997; Delmonte et al., 2004a) and interglacials (white open squares, this work) and comparison with the signature of samples from the Southern Hemisphere potential source area (PSA). The PSA samples were selected for their < 5 m size fraction (with the exception of Australian samples, selected for their : higher, Betula)

III

Betula–Pinus time (Betula Pinus)

II

Betula time (Betula > Pinus)

I

Afforesting time (Hippophae¨, Juniperus, Artemisia, Helianthemum, Thalictrum, tertiary sporomorphs)

Holsteinian in northern Germany. A main reason for this may be seen in edaphic conditions. In northern Germany, poor acid soils (podsols) ground on meltwater sands or tills, whereas loesses and brownearth soils are widespread in the Eifel and may be assumed also for pre-Holocene interglacials. On the other hand, higher values for Abies at mountain range sites in comparison with lowland sites have also been documented for the Eemian (Gru¨ger, 1995; Mu¨ller, 2000). The Betula–Pinus–Picea–Poaceae peak (Do¨ttingen LPAZ 4a) at the end of the Taxus period and the start of the Carpinus–Abies time allows comparison with the first Betula–Pinus peak from Munster/Breloh (Mu¨ller’s PAZ VIII). This likely equivalent to Mu¨ller’s older BPP is just about 5 cm in sediment thickness and was only detected through systematic searching of the relevant

core section. The peak in Poaceae pollen grains is only within sample 12.73–12.72 m, the Betula peak only within the following sample 12.72–12.71 m, whereas the Pinus peak occurs between 12.74 m and 12.69 m depth, centred at 12.73–12.71 m depth. The small depth range of this event shows that insufficient sampling resolution (compared to temporal resolution of the records) or even little reworking of the concerning sediment can conceal this feature. The observed ash layer (at 12.90–12.77 m depth) immediately below the LPAZ 3/4 boundary may explain the change in plant associations. In this case, the Betula–Pinus–Picea–Poaceae peak would represent pioneer vegetation after a catastrophic volcanic event. However, no charcoal or wood remains were detected, indicating that the location was not in close proximity to the corresponding

A New Holsteinian Pollen Record

413

Fig. 27.10 Simplified pollen profile from site Schmerz-Gro¨bern [originally in Eissmann et al., (1995), revised and redrawn after Ku¨hl and Litt, this volume] with pollen assemblage zones (PAZ) numbered after the classification of Erd. The setback of Carpinus after its initial expansion shall coincide with the Do¨ttingen LPAZ 4c and therefore with Mu¨ller’s second Betula–Pinus peak.

eruption. Beneath the ash layer, down to about 13.05 m depth, spanning the whole LPAZ 3b, the sediment layering is inclined (with ash remains at .075–13.055 m depth), indicating displacement likely in causal connection with the ash (volcanic tremor, earthquake, slump or ash load). This could also explain the conspicuous decrease in pollen concentration in that depth (Fig. 27.6). A destructive, disastrous temperature backstroke as a reason for the Betula–Pinus– Picea–Poaceae peak [as suggested by Mu¨ller (1974b)] for his first BBP seems unlikely, because thermophilous taxa like Quercus (and in general the mixed oak forest members) are unaffected. No significant signal can be obtained from cold-resistant, heliophilous herbs (e.g. Artemisa and Thalictrum), which would be expected if this event had a climatic cause. Furthermore, from the pollen concentration diagram (Fig. 27.7) it can be obtained that the Betula, Pinus and Picea

peaks in the pollen percentage diagram are not caused by extraordinary increases of the absolute pollen number of these taxa. In fact, their percentage peaks are mainly caused by a not proportionally steep drop of especially Corylus pollen concentration per volume. Only the Poaceae pollen percentage peak is accompanied by a steep increase in absolute pollen number. The presence of the Betula–Pinus peak both in the Eifel and in north Germany strengthens the demand for an over-regional connection. However, there are also differences between both sites. At Munster/Breloh, the event is characterized by a Betula–Pinus peak in combination with steep decreases of Alnus and Corylus, which both recover afterwards. In contrast to Munster/Breloh, the event at Do¨ttingen marks the beginning decrease of Corylus which does not show a (re)expansion after the event, even to higher percentages than before as at Munster/Breloh. Within the

414

Markus Diehl and Frank Sirocko

Do¨ttingen sequence, the Betula–Pinus–Picea– Poaceae peak is immediately followed by peaks of Corylus and Quercus. At Do¨ttingen, no setback of Alnus was found. The Do¨ttingen core reflects another feature that is also evident in several Holsteinian sites (Munster/Breloh, Klieken and Gro¨bern) and which might be correlated between these profiles. This is the contemporarily setback of Carpinus in LPAZ 4c after its initial expansion. At Munster/ Breloh, this event coincides with the second Betula–Pinus peak (Mu¨ller’s PAZ XI), at Gro¨bern with a small Pinus peak (between 19.00 and 19.50 m depth). The presence of thermophilous taxa (e.g. Tilia, Vitis, Buxus, Osmunda etc.) at Do¨ttingen seems to contradict a connection of this phase with climate deterioration, in particular because no signal comes Table 27.4 Tentative pollen zone correlation between Mu¨ller´s and Erd´s pollen assemblage zones (PAZ), and the local pollen assemblage zones (LPAZ) from Do¨ttingen (BPP: Betula– Pinus peak) PAZ After Mu¨ller

PAZ After Erd

¨3 LPAZ DO

XV

Fuhnian

7b

XIV

7

7a

XIII

5, 6 C

XII

6

4d

5

4c

B A (2. BPP)

XI (2. BPP) X IX

4b

VIII (1. BPP)

4

C VII

B

4a 3b

3

3a

2

2

A VI V IV

1b

III II I

1 1a

from Pinus, Betula or from the nonarboreal taxa. Instead, an ash stripe of about 2 cm thickness lies within this zone and may point again to an interaction with vegetation dynamics. ¨ 3 PAZ with A tentative correlation of the DO Erd’s and Mu¨ller’s palynological classification of the Holsteinian is done in Table 27.4. From the pollen zone correlation presented in Table 27.4, Mu¨ller’s second Betula–Pinus peak was at the same time as the Carpinus setback in Do¨ttingen’s LPAZ 4c, whereas Mu¨ller’s older BPP coincides with Do¨ttingen’s LPAZ 4a, respectively the Taxus decline in the beginning of Erd’s PAZ 4. Do¨ttingen’s LPAZ 6 must have been of comparatively short duration but with high sedimentation rates, indicated by its sediment thickness and detritus content. 27.7 CONCLUSION (1) The Do¨ttingen pollen sequence belongs to the Holsteinian interglacial. It represents the vegetation succession for the low mountain range site of the Eifel. (2) The Do¨ttingen profile offers a bipartite sequence, split into a limnic earlier and a telmatic later part with increasing nonarboreal pollen percentages. An episode of peat formation marks the boundary of this bipartition. (3) The older Betula–Pinus peak from Munster/Breloh can also be observed at Do¨ttingen. A volcanic eruption occurred immediately before this event. No signal for climate deterioration during the corresponding Do¨ttingen LPAZ 4a was found. (4) The Do¨ttingen profile documents a setback of Carpinus, correlatable with the younger Betula–Pinus peak from Munster/Breloh. No evidence for a causal connection to climate deterioration was found. Again a small ash stripe points to volcanic activity during this period. (5) The interglacial ends with an open boreal forest dominated by Betula, Pinus and Picea.

A New Holsteinian Pollen Record

In summary, the Holsteinian Do¨ttingen site reveals similar vegetation ‘anomalies’ as in northern Germany, but the Eifel record does not corroborate the existence of regionwide severe cold events within the interglacial. In contrast, the Eifel vegetation ‘anomalies’ develop subsequent or contemporarily to phases of volcanic activity. If this can be indeed interpreted as a climate/plate tectonic relation or as a response of central European vegetation to volcanic activity in the Eifel remains to be discovered.

ACKNOWLEDGEMENTS The study was funded by the German Ministery for Research and Education in the frame of the DEKLIM program. Discussions with Charles Turner, Helmut Mu¨ller and Thomas Litt provided insights into the Holsteinian vegetation dynamics. We thank in particular Thomas Litt, who introduced the first author into the technique of pollen preparation and the taxonomy of palynomorphs.

REFERENCES Averdieck, F.R., 1992. Das Holstein-Interglazial von Hamburg-Hummelsbu¨ttel. Meyniana 44, 1–13. Berglund, B.E., and Ralska-Jasiewiczowa., 1986. Pollen analysis and pollen diagrams. In: B.E. Berglund (Ed.), Handbook of Holocene palaeoecology and palaeohydrology, pp. 455–484. John Wiley, Chichester. Beug, H.-J., 2004. Leitfaden der Pollenbestimmung fu¨r Mitteleuropa und angrenzende Gebiete. Friedrich Pfeil Verlag, Mu¨nchen. Bu¨chel, G., 1984. Die Maare im Vulkanfeld der Westeifel, ihr geophysikalischer Nachweis, ihr Alter und ihre Beziehung zur Tektonik der Erdkruste. Unpublished PhD Thesis, Johannes GutenbergUniversita¨t. Bu¨chel, G., and Lorenz, V., 1982. Zum Alter des Maarvulkanismus der Westeifel. Neues Jahrbuch fu¨r Geologie und Palaöntologie 163, 1–22. Bu¨chel, G., Negendank, J.F.W., Wuttke, M., and Viereck-Go¨tte, L., 2000. Quarta¨re und tertia¨re Maare der Eifel, Enspel (Westerwald) und Laacher See: Vulkanologie, Sedimentologie und Hydrogeologie.

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In: F.O. Neuffer, and H. Lutz (Eds.), Internationale Maar-Tagung. Mainzer Naturwissenschaftliches Archiv, Daun/Vulkaneifel. Cepek, A.G., and Erd, K., 1975. Das Holstein-Interglazial im Raum Neuruppin - ein neues pollenstratigraphisches Richtprofil und seine quarta¨rgeologische Bedeutung. Zeitschrift fu¨r geologische Wissenschaften 3, 1151–1178. Dassow, W., 1987. Neue Holstein-Interglazial-Profile aus dem Quarta¨r im Raum Leipzig. Zeitschrift fu¨r geologische Wissenschaften 15, 195–203. Eissmann, L., Litt, T., and Wansa, S., 1995. Elsterian and Saalian deposits in their type area in central Germany. In: J. Ehlers, S. Kozarski, and P.L. Gibbard (Eds.), Glacial deposits in North-East Europe. A.A. Balkema/Rotterdam/Brookfield, pp. 439–464. Erd, K., 1969. Das Holstein-Interglazial von Granzin bei Hagenow (Su¨dwestmecklenburg). Geologie 18, 590–599. Erd, K., 1970. Pollen-analytical classification of the middle Pleistocene. In: The German Democratic Republic. Palaeogeography, Palaeoclimatology, Palaeoecology 8, 129–145. Erd, K., 1973. Vegetationsentwicklung und Biostratigraphie der Do¨mnitz-Warmzeit (Fuhne/Saale 1) im Profil von Pritzwalk/Prignitz. Abhandlungen des Zentralen Geologischen Instituts 18, 9–48. Erd, K., 1987. Holsteininterglaziale Ablagerungen von Rossendorf bei Dresden. Zeitschrift fu¨r geologische Wissenschaften 15, 281–295. Erd, K., and Mu¨ller, A., 1977. Die Pleistoza¨nprofile Prellheide und Wildschu¨tz, Bezirk Leipzig, mit vollsta¨ndigem Holstein-Interglazial. Zeitschrift fu¨r geologische Wissenschaften 5, 745–765. Erdtman, G., 1969. Handbook of Palynology: An Introduction to the Study of Pollen Grains and Spores. Munksgaard, Copenhagen. Faegri, K., and Iversen, J., 1989. Textbook of Pollen Analysis. Wiley, New York. Frenzel, B., 1968. Grundzu¨ge der pleistoza¨nen Vegetationsgeschichte Nord-Eurasiens. Franz Steiner Verlag, Wiesbaden. Gistl, R., 1928. Die letzte Interglazialzeit der Lu¨neburger Heide pollenanalytisch betrachtet. Botanisches Archiv 21, 648–710. Gru¨ger, E., 1995. Correlation of Middle-European Late-Pleistocene pollen sequences of the Pfefferbichl and Zeifen types. Mededelingen Rijks Geologische Dienst 52, 97–104. Hallik, R., 1960. Die Vegetationsentwicklung der Holstein-Warmzeit in Nordwestdeutschland und die Altersstellung der Kieselgurlager der su¨dlichen Lu¨neburger Heide. Mitteilung aus dem Geologischen Landesamt Hamburg 32, 326–333. Jerz, H., and Linke, G., 1987. Arbeitsergebnisse der Subkommission fu¨r Europaïsche Quarta¨rstratigraphie: Typusregion des Holstein-Interglazials

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(Berichte der SEQS 8). Eiszeitalter und Gegenwart 37, 145–148. Linke, G., and Hallik, R., 1993. Die pollenanalytischen Ergebnisse der Bohrungen Hamburg-Dockenhuden (qho 4), Wedel (qho 2) und Hamburg-Billbrook. Geologisches Jahrbuch A 138, 169–184. Litt, T., 1994. Palaöo¨kologie, Palaöbotanik und Stratigraphie des Jungquarta¨rs im nordmitteleuropaïschen Tiefland. Dissertationes Botanicae 227. Litt, T., 1999. Bio- and chronostratigraphy of the lateglacial in the Eifel region, Germany. Quaternary International 61, 5–16. Litt, T., 2000. Vegetation history and palaeoclimatology of the Eifel region as inferred from palaeobotanical studies of annually laminated lake sediments. Terra Nostra 6: International Maar Conference, Daun/Vulkaneifel, 259–263. Lorenz, V., and Zimanowski, B., 2000. Vulkanologie der Maare der Westeifel. In: F.O. Neuffer, and H. Lutz (Eds.), Internationale Maar-Tagung. Mainzer Naturwissenschaftliches Archiv, Daun/Vulkaneifel. Majewski, J., 1961. Pollenanalytische Untersuchung der Kieselgur von Klieken. Geologie 32, 10–14. Menke, B., and Tynni, R., 1984. Das Eeminterglazial und das Weichselfru¨hglazial von Rederstall/Dithmarschen und ihre Bedeutung fu¨r die mitteleuropaïsche Jungpleistoza¨n-Gliederung. Geologisches Jahrbuch 76. Meyer, K.-J., 1974. Pollenanalytische Untersuchungen und Jahresschichtenza¨hlungen an der holstein-zeitlichen Kieselgur von Hetendorf. Geologisches Jahrbuch A 21, 87–105. Moore, P. D., Webb, J. A., and Collinson, M. E., 1991. Pollen Analysis. Blackwell Scientific Publications, Oxford. Mu¨ller, H., 1974a. Pollenanalytische Untersuchungen und Jahresschichtenza¨hlungen an der eem-zeitlichen Kieselgur von Bispingen/Luhe. Geologisches Jahrbuch A 21, 149–169. Mu¨ller, H., 1974b. Pollenanalytische Untersuchungen und Jahresschichtenza¨hlungen an der holstein-zeitlichen Kieselgur von Munster-Breloh. Geologisches Jahrbuch A 21, 107–140. Mu¨ller, U. C., 2000. A Late-Pleistocene pollen sequence from the Jammertal, south-western Germany with particular reference to location and altitude as factors determining Eemian forest composition. Vegetation History and Archaeobotany 9, 125–131.

Mu¨ller, H., and Ho¨fle, H.-C., 1994. Die HolsteinInterglazialvorkommen bei Bossel westlich von Stade und Wanho¨den no¨rdlich Bremerhaven. Geologisches Jahrbuch 134, 71–116. Neumann, F., 2000. Pollenanalyse des Holsteinvorkommens in Klieken und palynostratigraphische Anwendung am Ostufer des Arendsees. Unpublished Diploma Thesis, Rheinische Friedrich Wilhelm-Universita¨t Bonn. Pirrung, B. M., 1998. Zur Entstehung isolierter alttertia¨rer Seesedimente in zentraleuropaïschen Vulkanfeldern. Mainzer Naturwissenschaftliches Archiv 20. Schmincke, H. U., 2000. Vulkanismus. Wissenschaftliche Buchgesellschaft Darmstadt. Selle, W., 1954. Die Vegetationsentwicklung des Interglazials von Ober-Ohe in der Lu¨neburger Heide. Abhandlungen vom Naturwissenschaftlichen Verein zu Bremen 33, 457–463. Selle, W., 1955. Die Vegetationsentwicklung des Interglazials vom Typ Ober-Ohe. Abhandlungen vom Naturwissenschaftlichen Verein zu Bremen 34, 33–46. Sirocko, F., Seelos, K., Schaber, K., Rein, B., Dreher, F., Diehl, M., Lehne, R., Ja¨ger, K., Krbetschek, M., and Degering, D., 2005. A Late Eemian Aridity Pulse in central Europe during the last glacial inception. Nature 436, 833–836. Stachel, T., and Bu¨chel, G., 1989. Das Do¨ttinger Maar: Fallstudie eines großen tertia¨ren (?) Tuffschlotes im Vulkanfeld der Hocheifel. Zeitschrift der deutschen geologischen Gesellschaft 140, 35–51. Stockmarr, J., 1971. Tablets with spores used in absolute pollen analysis. Pollen et Spores 13, 615–621. Straka, H., 1957. Pollenanalyse und Vegetationsgeschichte. A. Ziemsen-Verlag - Wittenberg Lutherstadt, Kiel. Straka, H., 1975. Pollen und Sporenkunde: Eine Einfu¨hrung in die Palynologie. In: Grundbegriffe der modernen Biologie. Fischer Verlag, Stuttgart, pp. 238. Szafer, W., 1953. Stratygrafia plejstocenu w Polsce na podstawie florystyczneiy (Pleistocene stratigraphy of Poland from the floristical point of view). Rocz. Pol. Tow. Geol. 22, 1–99. van den Bogaard, B., and Schmincke, H.-U., 1990. Die Entwicklungsgeschichte des Mittelrheinraumes und die Eruptionsgeschichte des Osteifel-Vulkanfeldes. In: W. Schirmer (Ed.), ‘‘Rheingeschichte zwischen Mosel und Maas.’’, Hannover, pp. 166–190.

28. Interglacial Pollen Records from Scho¨ningen, North Germany Brigitte Urban University of Lu¨neburg, Campus Suderburg, Herbert-Meyer-Str. 7, 29556 Suderburg, Germany

ABSTRACT

28.1 INTRODUCTION

The Pleistocene sequence of the Scho¨ningen lignite mine contains a number of interglacial and interstadial limnic and peat deposits, travertine tuff, soils, tills and fluvioglacial as well as loess deposits. There is evidence of four interglacials younger than the Elsterian glaciation and preceding the Holocene. The complex Pleistocene sequence contains six major cycles. The sequence begins with Late Elsterian glacial and interstadial deposits preceding the Holsteinian, followed by the Reinsdorf and Scho¨ningen interglacials, which represent the pre-Drenthe (Early Saalian Stadial) period. A pedocomplex developed in alluvial loess, younger than the Drenthe Stadial of the Saalian glaciation, is succeeded by a sequence of soft travertine and peat representing Eemian marine isotope stage 5e (MIS 5e) and MIS substages 5d, 5c. Channel sediments provide evidence of the Weichselian late glacial and the Holocene. The Scho¨ningen pollen record of MIS 5 and tentative correlatives of MIS 7, 9 and 11 as well as temperate interstadials of Late Elsterian and (intra) Saalian (s.l.) age are discussed and compared with other pollen records of North Germany, Western and Central Europe. Detailed pollen evidence for the Reinsdorf sequence, which is significant for its Lower Palaeolithic sites, including stadials and interstadials, is a major focus.

The Pleistocene sequence of Scho¨ningen (Lower Saxony, Germany) provides a key link between unglaciated and glaciated areas in Western Central Europe (Fig. 28.1). The complex Pleistocene sequence is of significance for the subdivision of the glaciated younger Middle Pleistocene part of Western Central Europe (Thieme et al., 1987; Urban et al., 1988, 1991a, 1991b; Urban 1995a, 1995b, 1996a, 1996b, 1999, 2002) and for archaeological evidence of early human occupation by Homo erectus (Thieme et al., 1987, 1992, 1993; Thieme and Maier, 1995). The investigations have focussed on exposed Quaternary deposits that are considered to span much of the last 500 000 years. The investigations occurred as excavation fronts progressed during mining of the underlying Eocene lignites. The Quaternary deposits are composed of various types of sediments including peaty, muddy and silty layers from former swamps, lakes, peat bogs and river flood plains, as well as hydromorphic soils that contain characteristic pollen assemblages. Fossil remains of molluscs, small and large mammals, fishes, reptiles and plant macro fossils are fairly abundant in some layers. The classical Holsteinian interglacial underlain by Elsterian glacial and Late Elsterian interstadial sediments was followed by a cold climatic deterioration interrupted by several temperate phases, the Mißaue 1, Mißaue 2 and SU A interstadial (Urban et al., 1991b). This cycle was later termed Channel 1 (Fig. 28.2) (Mania, 1998). A series of six such cycles of interglacial or

Keywords: Pleistocene, palaeobotany, palaeoclimate, 230Th/234U dating, interglacials

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and Channel III (Scho¨ningen Interglacial) represent warm climatic periods older than the Saalian ice advance. Channel IV, a pedosequence of pseudogleyic, alluvial loess, which has not yet been fully studied (Altermann, in preparation), overlies the older Saalian till (Drenthe stadium). The sediment in Channel V is a sequence of travertine, silts and peat, which has been correlated to the Eemian and to stage 5e of the marine isotope record (Urban et al., 1991a; Heijnis, 1992; Heijnis and Urban, 1995). Early Weichselian silts composed of loess, solifluction layers and fluvial deposits mark the onset of strong cooling. Late Weichselian Allero¨d peat with the Laacher See volcanic tuff layer and Younger Dryas silt had been identified underlying the Holocene sequence (Channel VI) (Urban et al., 1988) (Fig. 28.2). There is still debate about the stratigraphic position and correlation of the Reinsdorf and the Scho¨ningen interglacials and major interstadials found in Scho¨ningen with other pollen records and the marine

1 4

3 2

Weichselian Saalian 1- Wacken 2- Schöningen 3- Pritzwalk / Prignitz 4- Hoogeveen

Fig. 28.1 Maximum extension of Saalian and Weichselian ice sheets. Locations of Hoogeveen, Wacken, Scho¨ningen and Pritzwalk / Prignitz (Do¨mnitz).

interstadial and early glacial deposits in depressions (channels) have been identified, all considered to be climatically induced. Channel II (Reinsdorf Interglacial)

Exposure at the open-cast lignite mine, Schöningen NE

SW Esbeck Elm

Rim syncline

Salt dome

I II

III

IV

V

VI

b

1

4

2

5

3

6

~500 m

7

10

8

11

9

b

12

Fig. 28.2 Schematic section through the Quaternary sedimentary cycles I–VI (modified after Mania 1995; Mania, 1998; Thieme, 1997). The thickness of the geological deposits is not shown to scale. The Quaternary sediments reach a maximum thickness of ca. 45 m in the Esbeck mining field (Urban et al., 1991b). The actual distance between cycle VI Channel filling and the salt dome is about 2 km (this distance is not shown to scale). 1, Elsterian glacial deposits; 2, Saalian glacial deposits; 3, lacustrine deposits; 4, limnic telmatic sequences; 5, soil complexes; 6, loess deposits; 7, evaporites; 8 gypsum cap rock; 9, Buntsandstein; 10, Triassic limestone (Muschelkalk); 11, Triassic deposits (Keuper); 12 ð¼ bÞ, Tertiary deposits.

Interglacial Pollen Records from Scho¨ningen

istotope statigraphy. Of particular interest is the age and stratigraphic position of the Reinsdorf sequence which contains archaeological horizons with wooden throwing spears that are thought to be the oldest hunting weapons so far discovered (Thieme, 1996, 1997, 1998, 1999). 28.2 OPEN MINE ESBECK/ ¨ NINGEN SCHO The halokinetic depression of Scho¨ningen formed on the flanks of the HelmstedtStaßfurt Zechsteinian salt dome and was filled with thick limnic and marine sequences during Eocene times. At Scho¨ningen, these Tertiary strata are unconformably overlain by Quaternary sediments and soils of Middle and Late Pleistocene and Holocene age (Fig. 28.2). Mania (1998) described the Quaternary halokinetic processes as persisting mainly during interglacial periods, when shallow depressions, the channels, were formed, trending predominantly from northwest to southeast, corresponding to the trend of the salt dome. These channels have been filled with thick limnic, telmatic sediments or pedocomplexes (Fig. 28.2). Fennoscandian glaciers advancing from the northeast filled the depressions with glacial deposits. The study area has been covered twice by a Fennoscandian ice sheet, first during the Elsterian, the oldest ice advance of Northern Germany and Central Europe, and then the later Saalian advance (Fig. 28.1), and it was unglaciated during the last glaciation (Weichselian, MIS 4-2). 28.2.1 Elsterian The oldest Pleistocene sediments are glaciofluvial sands of an early Elsterian ice advance, which are overlain by two Elsterian tills. The younger till is capped by rhythmites, which in turn are overlain by three interstadial layers, formed in shallow basins, locally named Offleben 1, Offleben 2 and

419

Esbeck (Urban et al., 1988, 1991b; Elsner, 2003; Wansa, unpublished data) (Fig. 28.3). During the Offleben 1 and Offleben 2 interstadials, a boreal forest type prevailed, mainly dominated by birch, pine and spruce and locally elder in the late phases of a growing fen. The youngest of the investigated interstadials, the Esbeck Interstadial (Fig. 28.4), which is documented by nearly 3 m of organic mud and peat, is characterised by an early phase of open tundra-steppe with Juniperus, Salix, Pinus, Betula and some Alnus and Picea, followed by a succession of Juniperus and Salix during the Betula Zone (ESB3) and the expansion of Pinus and Picea during the local fen growth Zone ESB4, which is characterised by the occurence of Larix pointing to continental climatic conditions. Possibly three lithologically corresponding Late Elsterian interstadial phases have recently been identified in core sediments taken in the neighbouring Aller valley near Morsleben (Saxony-Anhalt), an exploration site for nuclear waste disposal, by Elsner (2003). 28.2.2 Holsteinian (Cycle I) During the initial phase of the brown coal excavations, limnic and telmatic deposits of a channel (Cycle I, Fig. 28.2) overlying the Elsterian sequence have been exposed adjacent to the Eemian Channel (Cycle V) and stratigraphically well positioned in relation to the Holocene (channel and Cycle VI) (Fig. 28.5). Because of its palynological composition in comparison with other sites in NW Germany (amongst others Munster-Breloh: Mu¨ller, 1974b; Hamburg-Dockenhuden: Linke and Hallik, 1993; Bossel: Mu¨ller and Ho¨fle, 1994), Urban et al. (1991b) assigned Cycle I to late parts of the Holsteinian interglacial. The terminal phase of the Holsteinian (Fig. 28.3) contains Abies, Pinus, Picea and Pterocarya as well as the water fern Azolla filiculoides, which is restricted to Early and Middle Pleistocene interglacials in

420

MITTELPLEISTOZÄN (SPÄT-) ELSTER OFFLEBEN-INTERSTADIAL I

STA- OFFLEBEN DIAL INTER?

ESBECK-STADIAL

STAD II

SAALE i. w. S.

HOLSTEIN

ESBECKINTERSTADIAL

BUSCH HAUS A

MISSAUE I I/II II

BUSCHHAUS B

INTERSTADIAL SU A

STADIAL SU A

HIATUS

Brigitte Urban

OF4/5 OF 1

OF 2

OF3

ESB 3 OFII1

ESB 1

ESB 2

ESB 4 ESB 5

Fig. 28.3 Elsterian and Holsteinian pollen sum curves.

SU8 SU9 SU10

SU 3 SU 1

SU 2

SU 4

SU 5

SU6 SU7

SU 11

SU 12

SU 13


Fig. 28.4 Limnic telmatic sediments of Late Elsterian Esbeck interstadial overlaying silt and niveofluviatile sand.

Central Europe (Urban, 1997). A strong cooling is recorded by a significant increase of Artemisia and grasses during the following Buschhaus A Stadial, which is considered to mark the onset of the Saalian Complex sensu lato (penultimate glacial complex) (INQUA SEQS, 1992). That stadial period is followed by a twofold temperate phase, the Missaue I and II interstadials, characterized by pine, birch and some spruce. Buschhaus B Stadial had a steppe environment with dwarf shrub tundra and was followed by another temperate phase, interstadial SU A, with Pinus being the dominant tree genus (Fig. 28.3). Cycle I sediments are not recorded from the southern mining area of the Scho¨ningen open-cut mine, but here evidence for a Late Elsterian fire place of Homo erectus has been found, which is situated discordantly below gravel and Channel II (Cycle II) sediments. Richter (1998) dated a burned flint of the fire place by thermoluminescence (TL) to 450 40 kyr.

421

Fig. 28.5 Left (position of mining device): Cycle VI (late glacial and Holocene deposits composed of peat, alluvial loess, Tschernozem). Middle Part: Cycle V (Eemian calcareous tuff and peat and Early Weichselian loess palaeosol sequence). Right: Channel I (Holsteinian and successive early Saalian interstadials: peat and silty basin deposits). Photo from Thieme and Maier, 1995 (modified).

28.2.3 Reinsdorf (Cycle II) Channel II (Figs. 28.2, 28.7) is filled by sediments of the Reinsdorf Interglacial, a new biostratigraphic unit between the Elsterian and the Saalian sensu strictu (Urban, 1995a) and further interstadials and stadials. The sediment sequence of this Cycle II contains a series of five levels (levels 1–5, Cycle II-1 to Cycle II-5) represented by peat and organic, silty and calcareous muds, in places extremely rich in molluscs (Mania, 1998). These lacustrine sediments of Cycle II have been found to occur at archaeological sites Scho¨ningen 12 and 13 (Thieme et al., 1993; Thieme and Mania, 1993; Thieme, 1996, 1997, 1999; Urban, 1999), marking the extension of the basin (channel), which had a width of at least 1000 m. Recent investigations give evidence for at least 13 local pollen assemblage zones (LPAZ). Eleven of them are presented in

422

JUNGPLEISTOZÄN SAALE i. e. S.

SAALE i. w. S. BUDDENSTEDTINTERSTADIAL I

ELM. BUDDENSTEDT- ELMSTAD B INTERSTADIAL II STAD C

DRENTHE/ WARTHE

WEICHSEL

EEM I

II

III

FRUH-

BIOSTRATIGRAPHIE BIOSTRATIGRAPHY

III

TIEFE/DEPTH (cm) LITHOLOGIE/LITHOLOGY THERMOPHILE GEHOLZE THERMOPHILOUS TREES ABIES FIR

BAUMPOLLEN ARBOREAL POLLEN

? POAGEAE GRASSES

ARTEMISIA WORMWOOD

UBRIGE TERRESTRAISCHE KRAUTER FURTHER TERRESTRIAL HEADS

CYPERACEAE

ERICACEAE

S1 S3 S2 S4 S5 S6 S7

S8

EA1

BI1

EB1

BII1

BII2

EC1

Fig. 28.6 Scho¨ningen and Eemian pollen sum curves.

T1a T1b T2a T2b

T3

T4

T5

T6

SFW1

LOCAL POLLEN ASSEMBLAGE ZONES

Brigitte Urban

ESBECK/SCHÖNINGEN

ELM-STADIAL A

SCHÖNINGEN


423

st ra

cl e cy en

ni ng

tig ra II-4c/5 S ph ch y ö

Middle Pleistocene

Saalian Complex s. l.

II-4b II-4a II-3

Reinsdorf stadial C Bio

Reinsdorf interstadial B Reinsdorf interstadial A

Reinsdorf stadial A

II-1

Reinsdorf interglacial

R 3a

R 3b

R 4/5

RS I1RS I2

0 5 10 15 20 25 30 35 40 45 50 55 60 65 70 0 5 10 15 20 25 30 35 40 45 50 55 60 65 70 75 80 85 90 95 100 105 110 115 120 125 130 135 140 145 150 155 160 165 170

Reinsdorf stadial B

RI II2 RI I1

RI I2 RS II RI II1

0 5 10 15 20 25 30 35 40 45 50 55 60 65 70 75 80 85 90 95 100 105 110 115 120 125 130 135 140 145 150 155 160 165 170 175 180 185 190

II-2

y Th Aberm i Ar es oph bo il r e ous al tr G p ra ol ees s le n Ar ses te m Fu is rth ia er C yp terr e er ac stri ea al he e Er rb ic s al es RS III1 LP AZ

lo g

th D ep 0 5 10 15 20 25 30 35 40 45 50

Li th o

(c

m )

Schöningen 13/94 profile 4 13/96 profiles 1, 2, 3

20

40

60

80

100

Silt

30 10 %

%

CaCO3

%

Peat

Loam

Molluscs

*Throwing spears

Fig. 28.7 Pollen sum curves of the Reinsdorf sequence.

this article on new pollen diagrams (Fig. 28.7, Figs. 28.9–28.12) that show a fivefold division of the middle and upper part of the interglacial and a sequence of

five climatic oscillations subdivided into eight LPAZ following the interglacial (Table 28.1). From the relative high values for grasses and herbs in the inferred forested

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Brigitte Urban

Table 28.1 Pollen zones of the Reinsdorf sequence (Cycle II, levels 1–5) Local pollen assemblage zones (LPAZ)

Characteristic floral elements

Biostratigraphy

Schöningen cycle

? RS III 1

? Herbs– Poaceae– Betula Pinus– (Picea– Betula, cf. Larix)

? Reinsdorf Stadial C

II- 5 II- 4c/II-5

RI II2 RI II1 RS II RI I2

Betula–Herbs (cf. Larix) Juniperus– Poaceae–Herbs (cf. Larix) Pinus– Poaceae–Herbs (cf. Larix)

RI I1 RS I 2

Pinus–(Piceae– Alnus)–Herbs Poaceae–Herbs Cyperaceae–

RS I 1

Salix–Calluna

R 4/5

Pinus– Ericaceae

II- 4b Reinsdorf Interstadial B * thro wing spears made from spruce

Reinsdorf Stadial B

II- 4a II- 3

Reinsdorf Interstadial A Reinsdorf Stadial A

II-2

II-1

R4 Carpinus– Picea–Abies

Reinsdorf Interglacial

R3/R2 Corylus–Alnus R1 Quercus– Fraxinus–Tilia Earl y phases are lacking

periods of the interglacial, a warm dry forest steppe climate can be deduced. Level II-1 represents the maximum temperature and terminal phases of the Reinsdorf Interglacial (Fig. 28.8). The vegetation succession, described by LPAZ (see as well Urban, 1995a), is as follows: a Quercetum mixtum phase, LPAZ R3a, followed by a Corylus–Alnus phase, LPAZ R3a/b, a Carpinus–Picea–Abies phase, LPAZ R3b, and a Pinus–Ericaceae zone, LPAZ R4/5

(Figs. 28.9, 28.10). The water fern Azolla filiculoides occurs frequently during the Quercetum mixtum phase. The interglacial ends with an opening of the pine-birch forest and a strong increase of grasses, terrestrial herbs and Ericales (Reinsdorf Stadial A: LPAZ RS I1 and RS I2, Figs. 28.9, 28.10). Lake shore sediments of the Reinsdorf interglacial level II-1 (Figs. 28.9, 28.10) contain abundant wood (Schoch, 1999), plant macrofossils (Jechorek, 2000), bones of


Fig. 28.8 Exposure of Cycle II limnic telmatic sediments overlying sand and gravel, sequence II-1: Reinsdorf Interglacial. 2004.

large mammals such as rhinoceros (Stephanorhinus kirchbergensis), straight tusk elephant (Palaeoloxodon antiquus), bear (Ursus spec.), horse (Equus mosbachensis), red deer (Cervus elaphus), deer (Capreolus capreolus) and the bovide Bos primigenius (Thieme et al., 1993; van Kolfschoten, 1995). Fossil remains of small and large rodents are also quite common in level II-1 as are remains of molluscs (Mania and Mai, 2001), birds, fish, reptiles (Bo¨hme, 2000), beetles and other insects. Numerous flint artifacts, wooden tools made of fir, and bones with cut marks found in level II-1 (Thieme, 1996, 1997, 1999) indicate early human occupation of the site during Lower Palaeolithic times. Level II-1 is characterised by Acer

425

campestre, Acer tataricum, Tilia cordata, Fraxinus excelsior, Prunus spinosa, Cornus sanguinea and Crategus monogyna, elements of slightly open deciduous forest, and by Abies alba, Taxus baccata, Carpinus betulus, Sorbus torminalis, Berberis vulgaris, Sambucus nigra, Cerasus avium, Lonicera xylosteum, elements of mesophilous mixed deciduous forests as demonstrated by karpological findings (Jechorek, 2000). The occurrence of the Pannonian floral element, Linum austriacum (Jechorek, 2000), points to 1.5–2 C higher annual temperatures compared to the present. The occurence of Zannichellia palustris indicates the presence of saline water, while occurrences of Potentilla anserina, Rumex maritimus and Chenopodium rubrum may point to slightly saline soil conditions. The presence of saline conditions is supported by chemical analysis of the sediments (Urban, unpublished data). Sediments of level II-2 are calcareous muds, which, in combination with a marked increase of grasses and a predominance of herbs and shrubs, denotes cool climatic conditions of Reinsdorf Stadial A (Figs. 28.9, 28.10; Table 28.1). Level II-3 sediments are organic mud and peat representing an interstadial period, named locally Reinsdorf Interstadial A (Figs. 28.10, 28.11). The vegetation was dominated by Pinus, with Betula, Alnus, and a few Picea (LPAZ RI I1 and RI I2). Trees indicative of a warm climate are almost absent. Level II-4 (Figs. 28.11–28.13) contains two stadials separated by an interstadial (Reinsdorf Stadials B and C and Reinsdorf Interstadial B), which comprise a transition into level II-5. Sediments are calcareous mud, organic mud and peat. The onset of climatic deterioration in the lower part of level II-4 is reflected by a dominance of herbs indicating a steppe environment (level II-4a, Reinsdorf Stadial B: LPAZ RS II). The upper part of level II-4 is characterised by a Pinus forest, with Betula (level II-4b, Reinsdorf Interstadial B, LPAZ: RI II1 Betula–Pinus zone and RI II2 Pinus–Betula

426

Schöningen 13/96 profile 1, x 668.00 m y 2.00 m Trees

Shrubs

Ericales

Grasses + Herbs pe

bs

r he

(m al th s s es tri p s rub cale ss res us e e a r i n D Tre Sh Er Gr Te Pi

ta e la /m eo jor c e a n p ty ype la m lia le t agotagoisia sum a t a re c an an tm tal Ce Se Pl Pl Ar To

AZ LP RS l2

99.35

s x ru a ies ari ipe ula us e L t n c n Pi Ab cf. Ju Be Al

pe ya ix ty ae lix m ar a us el a e m r e c h s t s e a u s s u c is o y c la s s u ce te etr na ini ra ea lu cu s in nu lu s elt er str bu gu a ulu ra rp ilia ory uer lmu alix cer raxi opu agu f. C . Pt f. O am ran yric um ede x rica rica mp allu acc ype oac C Q U S A F P F c cf c S F M H H lle E E E C V C P T Ca

99.26

RS l1

99.15 99.06 98.94

ra

st

o Bi

cle

y ph ra tig

cy

g

in

ön

en

h Sc

ll-2

L)

AS

ty

Reinsdorf stadial A

a di

98.55 98.45 98.35 R 3b

98.26 98.15 98.05

98 97.85 97.65 R 3a

97.55 97.45 97.36 97.26 20 40 60 80100

%

20 40 60 20 20 20 40 20 % % % %% % % %

20 20 20 40 % % % % % % % % % %% %% % %% %% % %%% % % % % %%% %

Fig. 28.9 Cycle II-1, II-2 (Reinsdorf Interglacial, Reinsdorf Stadial A, profile 1, Scho¨ningen 13/96).

500 1000 number

Brigitte Urban

98.65

ll-1

98.76


R 4/5

98.85

RS 12

102.52 102.27

RS 11

101.99 101.76

cy

hy

102.74

en ng ni hö

II-3 II-2

103.00

101.48 101.25

R 4/5

100.99 100.75 100.50 100.26


R11

103.25

II-1

ae e pe ce ty o ia cea eae p e ula ae ula m a a a y i c s al ag i ri ac -t e n ce nd su re ant tem cho ter ter iac nu sa lpe tal Ce Pl Ar Ci As As Ap Ra Ro Fi To

cle

Terrestrial Herbs tig ra p

Grasses

Sc

he

pe e pe ty ae ty e typ m um s pe a e m . s y s a s f u s u m c t u l s u r s ae c a u e a r a i u u ix a s ar ibu pe ulu x ta rylu us xin erc us a rpin us ltis yria ngu rica der mul cac Eric pet llun cin pera ce a c g i . a u ce ie . L . V ni p li tu n li x a y Fr Q Um Ti Ca Fa Ce O Fr M He Hu Er cf Em Ca Va Cy Po Pi Ab cf cf Ju Po Sa Be Co Al

tra

l s s ria le se st s ir ca ras erre inu E G T P

Ericaies

os

r

sh

Shrubs

Bi

h pt es De Tre 103.49

d an

Trees

Reinsdorf interstadial A

A

s rb

Reinsdorf stadial A

(m

s ub


) SL

LP AZ

Schöningen 13/96 profile 2 x 653.00m y -988.00m

99.98 99.75 20 40 60 80 100 %

20 40 60 80 % % % % % % % %

20 % %

20 % % % % %

20 % % % % % % % % % % % % %

20 %

20 40 % % % % % % % % % % %

500 1000 number

Fig. 28.10 Cycles II-1, II-2, II-3 (Reinsdorf Interglacial, Reindorf Stadial A, Reinsdorf Interstadial A, profile 2, Scho¨ningen 13/96).

427

428

Schöningen 13/96 profile 3, x 729,00 m y –993,00 m le

RS II

102.00 101.86

RS II

101.62 101.50

RI 2

101.43 101.37 101.25 101.11

RI 1

100.99 100.87 100.75

en ng ni hö Sc

100.62 100.50

20 40 60 80100 20 40 60 80

%

%

%% %%%%

20 40 20 % % %% % % % % %% % % % % % % % % %

20 % % % % % % %% % %% % %

20 40 % %%

500 1000 1500

number

Fig. 28.11 Cycles II-3, II-4a, II-4b (Reinsdorf Interstadial A, Reinsdorf Stadial B, Reinsdorf Interstadial B, profile 3, Scho¨ningen 13/96).

Brigitte Urban

101.74

ll-4b

102.10

ll-4a

102.20

ll-3

RS l2

102.30

cy c

hy tig ra p

102.35

tra

s x ru s a s ari pe ulu la ce bie f. L uni op alix etu i P A c J P S B

os

l s ia le ses str s ica as re u Er Gr Ter Pin

pe ty e e e . e ea ea p ae e pec a/m e yp ae ty pe ac c um p t ce y t e e m aty o s m ia odi ylla em m cea ae la e m . s s p c e ae dul um a ty tru a iu a f la s us ia ag go is op ph th tru ria ce s ae ncu u u a l c e e l u r u u e n n a n e n c m a a l a n c s s i i s u c re ant nt tem en ryo lia alic cho ter liu iac nu ry nus xin er u lia rp gu ltis ang ric m ica ica pe llu cc ype ac sa lipe tal Ce Pl Pla Ar Ch Ca He Th Ci As Ga Ap Ra Co Al Fra Qu Ulm Ti Ca Fa Ce Fr My Hu Er Er Em Ca Va C Po Ro Fi To

Bi

bs

r he

Reinsdorf interstadial B

d

an

bs

ru

sh

Terrestrial herbs

Reinsdorf stadial B

102.40

es

e Tr

A

Grasses

Reinsdorf interglacial A

h

pt

De

(m

Ericales

LP AZ

Trees and shrubs

) SL

20 40 60 80 % % % % % 20 % % % % % % % % 20 % 20 % % % % % % % % % % % % 20 % % %

II-4c/5

II-4b

40

Reinsdorf Stadial C

20

Reinsdorf Interstadial B

10


20 40 60 80 100 %

L)

e

cl

cy

y

ph

en

ng

ni

hö

Sc

ra

tig

tra

os

Bi

AZ

LP

C o Al rylu nu s C s ar Fr pin a u M ngu s yr la Er ica ic C ace al a l C una e yp er ac ea Po e ac e C ae er e Pl alia Ar ant typ te ag e C mis o m h ia a jo Th eno r/m a p ed C lictr odia ic u c ia ty As hor m ea pe e te iac C ra ea en ce e G tau ae al r R ium ea um Ap ex ia R ce an ae u R ncu os la Fi ace cea lip a e e To en ta du l s la um ty pe

Pi ce La a rix Ju n Sa ipe lix rus Be tu la

Te

Grasses

RS III1

sh

AS

Trees and shrubs

RI I2

d

an

m

(c

ru es bs rre s Pi t nu rial he s rb s

ss

ra

G

s

ee

Tr

th

ep

D

Schöningen 13/94 Profile 4, x 685.00m y 38.00m

Terrestrial herbs

1

50

200 400 number

Fig. 28.12 Cycles II-4b, II-4c/II-5 (Reinsdorf Interstadial B, Reinsdorf Stadial C, profile 4, Scho¨ningen 13/94). 429

430

Brigitte Urban

tartaricum, Acorellus pannonicus, Thymelaea passerina, Viola cf. alba, Ranunculus brutius, Ranunculus lateriflorus A. Dc. vl. nodiflorus L., Dychostylis micheliana and Potamogeton vaginatus. In agreement with the palynological findings, a warm climatic forest steppe type of vegetation indicating rather dry regional conditions can be inferred for the main part of the Reinsdorf Interglacial. Preliminary 230Th/234U age determinations of peat layers within the Reinsdorf interglacial deposits (level II-1) at site Scho¨ningen 12 (Heijnis and Urban, 1995) gave an approximate age of 320 kyr. 28.2.4 Scho¨ningen (Cycle III) Fig. 28.13 Cycle II, sequence II-4b: Reinsdorf interstadial B (strata of throwing spears) and sequence II-4c/II-5: Reinsdorf stadial C (cryoturbation of top sediments) 2004.

zone) (Figs. 28.12, 28.14). Pollen of Picea is very rarely found in Reinsdorf Interstadial B deposits. Jechorek (2000) identified macrofossils of Arctostaphylus uva-ursi, Carex aquatilis, Frangula alnus, Pinus sylvestris, Lonicera xylosteum, Rubus ideaus and Salix spec. in level II-4 sediments. Level II-4b also yielded numerous faunal remains, mainly horse, flint artifacts and well-preserved spears made from spruce, the oldest hunting spears so far discovered (Thieme, 1999). The pollen assemblage of level II-4b (Figs. 28.11, 28.12) contains only very rare Picea pollen, which might suggest transportation of the hunting weapons from scattered stands of spruce some distance from the site. Between level II-4c (Reinsdorf Stadial C, LPAZ: RS III1, nonArboreal zone) and II-5, within silty clays, frost structures occur and mark the onset of a periglacial environment and the definite end of cycle II (Fig. 28.13). In summary, the karpological investigations of the Reinsdorf sequence (Jechorek, 2000) reveal remains of 132 plant species including seven thermophilous exotics: Acer

Channel III contains a sedimentary sequence (Cycle III) that cross-cuts Channel II (Figs. 28.2, 28.6). The sequence is composed of silty muds and peats and represents the Scho¨ningen interglacial (Urban et al., 1991b; Urban, 1992, 1995a). The pollen assemblages are indicative of a warm, (sub)continental climate with high percentages of Pinus and Tilia with some Quercus. High components of Alnus found almost throughout the entire profile point to swampy environments. A Carpinus phase with Picea occurs near the end of the warm period. Abies is absent, apart from a single grain, while massulae of the water fern Azolla filiculoides are abundant in the Alnus-rich parts of the sequence. The Scho¨ningen interglacial is succeeded by the Elm A Stadial which is characterised by a marked increase of herbs and moderate increases of Artemisia, grasses and Ericales (Fig. 28.6). The Elm A Stadial is followed by two temperate periods, the Bu¨ddenstedt I and II Interstadials, with pollen assemblages indicative of Pinus–Betula forests. The Bu¨ddenstedt interstadials are separated by the Elm B Stadial. Increases of Betula and herbs mark the onset of another cold spell, the Elm C Stadial which caps the sequence (Fig. 28.9). The interglacial and stadial– interstadial sequence of Channel III (Cycle III) is overlain by glaciofluvial sands and till


Holocene Early

Middle

Weichselian

~30.000

Eemian

~115.000

~130.000

Warm desiduous forest

Boreal forest Steppe forest

Shrub tundra

Detailed subdivision of climatic 230 units of the upper Quaternary of Schöningen TH/U * cycle Central Europe. TL ** dating

Tentaive correlation

MIS

kyr

basically adapted from the Schöningen record

Subatianticum Subboreal Atlanticum Boreal Preboreal

Late

~10.000

Tundra

Years BP

Steppe

Palaeoecological evolution Polar desert

subdivision of the Quaternary

Glaciated landscape

Time scale

431

1 C VI

Late Dryas Bølling / Allenød Early Dryas

2

Denekamp Hengelo

3

Moershoofd Glinde Oerel Odderade

4 5a

Brörup

5c CV 115 – 149*

Eemian interglacial

5e

Late

Hiatus Warthe

C IV

Glaciation

6

Drenthe Hiatus Elm C Büddenstedt II Elm B

C III

Büddenstedt I Elm A

Schöningen interglacial Middle

Saalian Complex

~200.000

177 – 234*

7

prox. 320

7/9

Hiatus Level 5

Reinsdorf - Stadial C Reinsdorf - Interstadial B Reinsdorf - Stadial B Reinsdorf - InterStadial A

Level 4

C II

Reinsdorf - Stadial A

Reinsdorf - interglacial

Level 3 Level 2 Level 1

Hiatus Stadial SU A Interstadial SU A Buschhaus B

Early

Missaue I

Holsteinian

Missaue II Buschhaus A

CI

Holsteinian interglacial

Late

Hiatus Esbeck

9/11 450 ± 40** Burn Flint

Pleni-

Glacial

Elsterian

Offleben II Hiatus Offleben I

Glaciation

Fig. 28.14 Synthesis of the Scho¨ningen pollen records, dating and tentative correlation with the marine isotope stages.

of the first Saalian ice advance (Drenthe Stadium). Peat of the Scho¨ningen interglacial gave uncorrected 230Th/234U ages of 180 and 227 kyr (Heijnis, 1992). Based on

the pollen record, correlation has been made with the Wacken (Menke, 1980) and Do¨mnitz (Erd, 1973) interglacials (Urban et al., 1991b; Urban, 1995a).

432

Brigitte Urban

28.2.5 Cycle IV Channel IV, which is eroded into Saalian glacial deposits, contains a pedocomplex developed in alluvial loess. Two pseudogleyic layers, presently being analysed by Altermann (personal communication) suggest the presence of at least one periglacial phase between the two major ice advances of the Saalian glaciation (Drenthe and Warthe Stadium, Figs. 28.2, 28.14), which is recorded with tundra-type vegetation and soils from the Red Cliff on the Isle of Sylt (North Western Germany) (Felix-Henningsen and Urban, 1982) 28.2.6 Eemian (Cycle V) The sequence in Channel V (Cycle V) is represented by either a Luvisol developed in loess, or, due to the influence of the palaeodrainage system, a soft travertine and peat of last interglacial age. In Scho¨ningen, interglacial sedimentation started concurrently with a Carpinus phase and continued to the first Early Weichselian interstadial (Figs. 28.5, 28.6). The Eemian peat and travertine layers are rich in Abies during the Pinus-Picea-Abies-Zone. The travertine sediments were deposited during a period of about 6000 years, deduced from the reconstructed pollen zones (Mu¨ller, 1974a). Local hydrological conditions during the late last interglacial and early glacial periods have been determined by pollen analysis and by plant macro remains, specifically moss analyses (Ho¨lzer in Urban et al., 1991a). The Eemian peaty layers reveal a thorium/uranium age of 132 17 (Heijnis, 1992) (Table 28.2, Fig. 28.14).

vegetation history, palaeo landscape and the degree of human impact on the area (Figs. 28.2, 28.5) (Thieme et al., 1987; Thieme and Maier, 1995).

28.3 DISCUSSION The sequence of Scho¨ningen gives geological and palaeoecological evidence for several temperate phases between the Holsteinian and the Eemian and reveals data on three interglacials and at least 10 interstadials between the end of the Elsterian and the beginning of the Saalian (Drenthian) glaciation (s.str.). Some authors state that, as (temperate) deposits investigated in Scho¨ningen do not occur in the same outcrop in perfect superposition, the succession of the warm stages will remain to some extent debatable (e.g. Turner, 1998). In contrast, we believe, from intensive geological field work, mapping and analysis of long geological transects since the opening of mine Esbeck in the 1980s, the superposition of stratigraphic units, the overlapping of sedimentation units which cut or underlay the strata at certain times of excavation and adjacent occurrences of channel fillings, that a stratigraphic sequence has been well established (Elsner, 1987; 2003; Hartmann, 1988; Urban et al., 1988, 1991a, 1991b; Lenhart, 1989; Tschee, 1991; Thieme et al., 1993; Mania, 1995, 1998). Moreover, palaeoecological and biostratigraphic evidence, archaeological findings and radiometric dating have resulted in a comprehensive stratigraphic scheme for the glaciated margin of Western Central Europe.

28.2.7 Holocene (Cycle VI) The youngest sedimentation cycle in the Scho¨ningen mine is comprised of Late Weichselian and Holocene deposits and soils (Channel VI sediments), and braided river deposits of the Mißaue and its tributaries. The variety of both sediments and soils has allowed a detailed reconstruction of the

28.3.1 Late Elsterian, Offleben I and II and Esbeck Interstadials The Offleben I, II and the Esbeck interstadials, which are related to late phases of the Elsterian glaciation, but not to its late glacial period, occurring on top of the youngest Elsterian till, have no known biostratigraphic equivalent


433

Table 28.2 Tentative correlation of the Scho¨ningen sequence with the Velay record (Reille and de Beaulieu, 1995; Reille et al., 1998; de Beaulieu et al., 2001) Germany NE Niedersachsen

South Central France

Urban et al., 1988, 1991a, 1991b; Heijnis, 1992; Urban, 1995a, 1999, 2000; Urban and Heijnis, 1995

Eemian

230

Elm C Büddenstedt II Elm B Büddenstedt I Elm A Stadial Schöningen 230 Th/234U 180 and 227 kyr ? Harbke interstadial

Reinsdorf Stadial C

Le Bouchet 3 Belvezet Le Bouchet 2 Bonnefond Le Bouchet 1

(II-4/5) Charbonniers Stade (II-4)

Amergiers Interstade

(II-4)

Monteil Stade

Reinsdorf Interstadial A

(II-3)

Ussel Interstade

Reinsdorf Stadial A

(II-2)

Cayres Stade Landos

230

7?

? ?

Reinsdorf Stadial B Reinsdor f

5e

Ribains

Th/234U 132 +/- 17 kyr

Reinsdorf Interstadial B

MIS

Reille and de Beaullieu, 1995; Reille et al., 1998; de Beaullieu et al., 2001

Th/234U prox. 320 kyr?

40

Ar/39Ar 275 +/5kyr

Stadial SU A Bargette Stade Interstadial SU A Jagonas Interstade 2 Buschhaus B Stadial Pradelle 2 Missaue Interstadial (Missaue I, II) Jagonas Interstade 1 Buschhaus A Stadial Holsteinian Praclaux

elsewhere in Northern Germany and Western Europe. Attention should be paid to their relevance for correlation with the marine isotope record as they are preceding the classical Holsteinian interglacial and might be of significance for Late Elsterian climatic evolution. 28.3.2 Holsteinian, Missaue I and II Interstadials, Buschhaus B Stadial and Interstadial SU A In Scho¨ningen late phases of the Holsteinian interglacial and two following interstadials are documented (Urban et al., 1991a, 1991b; Urban, 1996a, 1996b). The terminal phases of the Holsteinian, which contain Abies, Pinus, Picea, Pterocarya and water fern Azolla filiculoides, have been

7/9 ?

9/11 ?

correlated with pollen zones XII, XIII and XIV (Mu¨ller, 1974b; Meyer, 1974). A strong cooling at the end of the Holsteinian marks the onset of the Saalian s.l.. This event, the Buschhaus A Stadial of the Scho¨ningen sequence, can be correlated with the Fuhne A Stadial at Pritzwalk/Prignitz (Erd, 1973). The twofold Mißaue interstadials I and II are equivalent to the Dockenhuden interstadial (Hallik and Linke, 1986; Urban, 1996a) and to the Pritzwalk interstadial A/B (Erd, 1973; Erd et al., 1987). In comparison with investigations at Bossel (northwest Niedersachsen), a site with sediments of the Holsteinian marine transgression (Mu¨ller and Ho¨fle, 1994), the early part of the twofold Mißaue interstadial may be correlated with zone XVI/XVII.

434

Brigitte Urban

Looking at the occurence, distribution and succession of certain taxa (Carpinus/Fagus, Pterocarya, Abies and Azolla filiculoides) in late Holsteinian sequences, not only of areas adjacent to the North Sea Basin (Zagwijn, 1973) but also of lower European latitudes, including the long sequence of the Velay (de Beaulieu et al., 2001) and La Cote in Vercors, France (Field et al., 2000), the sites of Samerberg II in Bavaria (Gru¨ger, 1983), Thalgut in Switzerland (Welten, 1988) and Krepiec, Zbojno and Losy in Poland (Lindner and Marciniak, 1998) are of significance. Long-distance correlation of the Praclaux interglacial with the Holsteinian has been proposed by de Beaulieu et al. (2001), and of the Mazovian by Lindner and Marciniak (1998), suggesting it a probable terrestrial correlative of stage MIS 11. Referring to Sarnthein et al. (1986), who 230 Th/234U and ESR dated marine molluscan shells from para-type and other Holsteinian interglacial deposits to > 350 and 370 kyr, Holsteinian beds have been correlated with MIS 11 or even an older interglacial event. There is ongoing debate on the exact age of the Holsteinian interglacial. Geyh and Mu¨ller (2005) recently presented 230Th/234U dates and a palynological review of the Holsteinian/Hoxnian interglacial. Their interpretation of 230Th/234U dates has led them to correlate the Holsteinian and the Hoxnian (Turner, 1970) with MIS 9. The authors present a brief review of correlations of the Holsteinian with MIS 11 citing mainly palynological work. Among those, their citation that, by using 230Th/234U dates of peat of the Scho¨ningen interglacial (Figs. 28.10, 28.12), Urban (1983, 1995a) ‘palynologically and indirectly correlates the Holsteinian interglacial to MIS 11’ is causing confusion and has to be corrected. As cited above, peat of the Scho¨ningen interglacial revealed uncorrected 230Th/234U ages of 180 and 227 kyr (Heijnis, 1992) and therefore can be related most appropriately to MIS 7 (Urban, 1995a). It can, therefore, not be used as an indirect correlation tool for dating the Holsteinian to MIS 11 (Fig. 28.14). There is further evidence that the Ka¨rlich interglacial sequence (Urban, 1983; Bittmann,

1992) might reveal two different interglacials (Urban, manuscript in preparation). Urban (1983), based on the palaeofloristic record, tentatively proposed a post-Holsteinian age for the Ka¨rlich interglacial s.str. Bittmann has correlated the Ka¨rlich interglacial sequence palynologically to the Bilshausen interglacial (Mu¨ller, 1965; Bittmann and Mu¨ller, 1996) and an 40Ar/39Ar age of about 400 kyr obtained from a tephra layer predating the Ka¨rlich interglacial (van den Boogaard et al., 1989: 396 þ 20 kyr) to the Cromerian (V). Refering to recent dating, van den Boogard (in Boenigk and Frechen, 1998) relates the eruption of the ‘Ka¨rlicher Brockentuff’ to the beginning of MIS 10. Boenigk and Frechen (1998) place the Ka¨rlich interglacial s.str. (Urban, 1983), by correlation with sequences from the Lower Rhine area, into the Saalian s.l., a correlation which had already been suggested by Urban (1983). The Holsteinian deposits of Scho¨ningen have not been dated so far. There is a TL date of a burnt silex from a prehistoric fire place (Richter, 1998) in Late Elsterian/Early Holsteinian deposits (Urban, unpublished data) available for the Scho¨ningen mine, revealing an age of 450 40 kyr (Fig. 28.14). The pollen record of those deposits point to an open tundra-(taiga) environment with pine and birch and indicate late glacial Elsterian environmental conditions. As there is dating and palaeoecological evidence of interglacial sequences following the Holsteinian, attention will be focussed on these superimposed strata of mine Scho¨ningen and their biostratigraphic correlation and preliminary dating. 28.3.3 Reinsdorf Interglacial, Reinsdorf Stadial A, Reinsdorf Interstadial A, Reinsdorf Stadial B, Reinsdorf Interstadial B, Reinsdorf Stadial C, Harbke Interstadial and Scho¨ningen Interglacial The term Reinsdorf Interglacial was introduced by Urban (1995a) for a new interglacial sequence at Scho¨ningen of post-Holsteinian and pre-Drenthe age (first


Saalian ice advance), which yielded abundant fossil remains as well as archaeological evidence for the presence and activities of Homo erectus (Thieme and Maier, 1995). The archaeological site is still exposed, and the excavation and research continues under the supervision of Hartmut Thieme (Archaeological Survey of Lower Saxony, Hanover). As stated earlier, the Reinsdorf sequence contains interglacial and stadial–interstadial floras (Table 28.1) that are quite distinct from the Holsteinian (Urban, 1995a, 1995b, 1999). The main characteristics of the Reinsdorf Interglacial are a climatic optimum characterised by a forest phase with the spread of Tilia before Corylus; which is only represented by low values, and by the occurrence of a late and less pronounced Abies phase. Furthermore, the Reinsdorf sequence is characterised by two pronounced interstadials interrupted by phases of climatic deterioration (stadials), when the vegetation opened up to grass and herb-rich steppic environments (Table 28.1, Figs. 28.9–28.12, 28.14). Compared to the Holsteinian vegetation reconstructed from the same outcrop, a warm and distinctive continental regional type of climate can be inferred for the Reinsdorf Interglacial, indicated, for example, by occurrences of Acer tartaricum and Larix, as well as by low representation of Corylus. Larix also occurs during the interstadial and stadial phases. These climatic interpretations are supported by the mollusc assemblages (Mania and Mai, 2001). A 16-m profile covering the biostratigraphic units of the Reinsdorf sequence at the excavation site is presently under investigation. Thermal ionization mass spectrometry (TIMS) 230Th/U dating of peat taken from this profile is currently in progress (Frechen et al., this volume). The Scho¨ningen interglacial fen peat deposits (Fig. 28.6) contain a distinctive vegetational succession, dominated by Pinus and Alnus throughout the entire thermal part of the sequence. The Scho¨ningen

435

Interglacial terminates abruptly and is succeeded by a taiga-tundra type of vegetation and then two short boreal conifer phases. Urban (1995a) already discussed the main differences between the pollen zones, marker species and local as well as regional environments in the Holsteinian, Reinsdorf and Scho¨ningen and the probable correlatives of the latter, the Wacken (Menke, 1980), Do¨mnitz (Erd, 1973) as well as the Eemian interglacial floras in the Scho¨ningen mine in great detail. It was concluded that the Reinsdorf and Scho¨ningen interglacials differ strongly from each other in the following features which have proven to be of biostratigraphic value for long-distance correlation: 1. Absence of Abies (Scho¨ningen), 2. Occurence of an Abies phase during the Carpinus–Picea phase (Reinsdorf) 3. Tilia peaks during the Corylus and early Carpinus phases (Scho¨ningen) 4. Expansion of Corylus with low values and only after that of the Mixed-oak forest phase (Quercetum mixtum, QM), dominated by Tilia (Reinsdorf) 5. Reinsdorf interglacial characterised by a warm, continental regional forest steppe climate 6. Dominance by Alnus and Pinus in all observed pollen zones (Scho¨ningen). 7. An abundance of Pinus but only local importance of Alnus reflecting the moist hydrological conditions of the stands (Reinsdorf) There are only few sites with records of similar age or of similar biostratigraphic significance to Scho¨ningen from neighbouring areas. However, the Holsteinian and postHolsteinian deposits from Morsleben and Ummendorf (Aller Valley, Saxony-Anhalt) described by Strahl (1999), which are located less than 10–15 km distance of the Scho¨ningen mine with comparable geological, geogenetical origins, are of special interest as they might contain equivalents of the Reinsdorf Interglacial. Strahl (1999) has

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considered the Aller interglacial, which follows the Holsteinian, the deposits of which are overlain by the Morsleben Stadial A, Morsleben Interstadial B and Morsleben Stadial C, a most probable time correlative of the Reinsdorf, Wacken (Menke, 1968) and Do¨mnitz in the profile of Pritzwalk/ Prignitz (Erd, 1973). As stated earlier, the Holsteinian at Scho¨ningen is followed by the Buschhaus A stadial, which is considered to mark the onset of the Saalian Complex s. l. This is followed by a twofold temperate phase, the Missaue I and II Interstadials and the Buschhaus B Stadial characterised by a steppe environment which was followed by another temperate phase, Interstadial SU A (Figs. 28.3, 28.14). Though in Morsleben only one interstadial (Morsleben Interstadial B) is recorded, which might point to some discordances/hiatus between the post Holsteinian interstadials and the deposits of the Aller interglacial in that section, it has certain similarities with the Reinsdorf interglacial. Strahl (1999) records four pollen zones of the Aller interglacial, including a Carpinus–Abies–Picea–Alnus zone, which have correlatives in the Reinsdorf interglacial pollen zones (R3–R5). Unfortunately, only descriptions of pollen zones without diagrams have been published by Strahl so far, prohibiting more detailed comparison. The correlation of the Aller interglacial and the Wacken and Do¨mnitz interglacials still remains rather tentative as they differ palaeobotanically strongly from the Aller and Reinsdorf interglacials. There is no evidence from the Morsleben site of two interglacials following the Holsteinian and preceding the first Saalian ice advance. However, two interglacials are recorded by Mu¨ller (personal communication, 1999 and 2005) from pit Nachtigall (Niedersachsen). Mu¨ller found evidence of a first interglacial, which is rich in Tilia and characterised by a late Abies phase immediately following the Holsteinian and the early Saalian interstadials described in Scho¨ningen (Urban, 1999), and of a

subsequent interglacial, which is lacking Abies and which is dominated by Alnus and Pinus. In summary, it is most likely that the earliest interglacial of pit Nachtigall is a correlative of the Reinsdorf Interglacial and the second one probably of the Scho¨ningen Interglacial. Another probable botanical equivalent might be found in the cores drilled in Go¨ttingen recorded by Gru¨ger et al. (1994). The diagrams b, c and d of Go¨ttingen Ottostraße are reflecting an interglacial sequence, which is interpreted by Gru¨ger et al. (1994) as representing three interglacials interrupted by stadials and followed by interstadials. The whole sequence is characterised by several hiatuses and other disturbances, which is reflected by ‘noise’ in the pollen record. If the diagram zones DA 1–17 (after Gru¨ger et al., 1994) are synthesised, taking into account the hiatuses and some probable sediment contamination including the solifluction layer, the initial phases of one interglacial and at least one following interstadial could be identified as having definite similarities with the Reinsdorf interglacial sequence. The pollen diagram of the lacustrine deposit of Elsterian to Saalian age (Behre, 2004) at Surheide near Bremerhaven spans the lower and middle parts of an interglacial which is characterized by a pronounced peak of Corylus, a late maximum of Abies and a lack of Carpinus. Behre consequently suggests correlation with pre-Holsteinian rather than Holsteinian or early Saalian sequences of intra-Elsterian age, focussing on the Ferdinandow interglacial (JancykKopikowa, 1975; Jancyk-Kopikowa and Zarski, 1995). The interglacial deposits of Surheide do not have a correlative in the Scho¨ningen sequence at the present state of investigation. At Ro¨persdorf (Erd, 1987), the terminal part of the recorded interglacial is almost lacking. However, based on the pattern of spread and dominance of Tilia and Corylus, ¨ cker interseveral authors correlated the U glacial from Ro¨persdorf with postHolsteinian and pre-Eemian interglacial


deposits. Based on recent results obtained within a drilling project of the Geological Survey of Brandenburg (Landesamt fu¨r Bergbau, Geologie und Rohstoffe Brandenburg) in the area of Prenzlau, Hermsdorf and Strahl (2005) point out major disturbances of the lithological setting of the interglacial deposits of Ro¨persdorf. Similar to Erd (1970), they correlate the interglacial deposits of Ro¨persdorf with the Eemian interglacial. Mania (1998) correlates the Reinsdorf sequence with Bilzingsleben II, which he considers to date to 400 kyr based on ESR and 230Th/234U determinations. Consequently, he relates it with MIS 11 and the Scho¨ningen Interglacial with Bilzingsleben III, which he correlates with the Do¨mnitz interglacial (Erd, 1973), citing an 230 Th/234U age of 320 kyr for the latter interglacial, which he therefore relates to MIS 9. As there are no comparable longer pollen records available for the travertine sequence of Bilzingsleben, a correlation with the Scho¨ningen sequence in our opinion is rather tentative. The Reinsdorf interglacial has been correlated by Urban (1995a) with the Zbojnian interglacial in Poland (Lindner and Marciniak, 1998), situated some 500 km east of Scho¨ningen. The Zbojnian interglacial is intercalated between two stadials that follow the Mazovian interglacial which is considered the equivalent of the Holsteinian in Poland. Correlation has been based on the great similarities of the climatic optimum characterised by Tilia dominance during the QM zone and the spread of Tilia before Corylus. Late, terminal phases of the Zbojnian interglacial are characterised by an Abies phase with lower values compared to the Mazovian interglacial, which is identical with the Abies distribution during the Holsteinian and Reinsdorf observed at Scho¨ningen. De Beaulieu et al. (2001) have correlated the Landos interglacial with the Zbojnian, ¨ cker (Erd, 1987) the Reinsdorf and the U Interglacials due to its stratigraphic position below the Bouchet 1 interglacial which is almost lacking Abies. Consequently, they

437

correlate Bouchet 1 with the Scho¨ningen, Wacken (Menke, 1968) and Do¨mnitz (Erd, 1970) interglacials. It should be noted that long-distance correlation based on vegetation changes and similarities inferred from pollen data and other plant remains between sites with latitudinal, altitudinal and edaphical differences have to take into account dissimilarities in occurrences of taxa or their relative representation (de Beaulieu et al., 2001). Considering these uncertainties, these authors state that marker taxa such as Pterocarya are very important. Fagus is also important in that it plays a key role in both the Holsteinian and in younger interglacials in the southern part of Central Europe and in contrast to Northern Germany where it appears with very low values during late phases of the Holsteinian, co-occuring with Pterocarya. With reservations and pointing out the tentative character of our correlation with the sequence of Velay (de Beaulieu et al., 2001) and La Coˆte, Val-de-Lans basin, France (Field et al., 2000), best-estimate comparison of Scho¨ningen with those long terrestrial records for the Middle Pleistocene of Central Europe is presented in Table 28.2. This palynological correlation is supported by the 40Ar/39Ar age of 275 5 kyr of the Armargier interstadial following the Landos interglacial that is tentatively correlated here with the Reinsdorf Interglacial. A recent reinvestigation of the Meikirch pollen record (Welten, 1982, 1988) in the Swiss Alpine Foreland by Preusser et al. (2005) has led to a reinterpretation of its correlation. Whereas Welten (1982, 1988) correlated the youngest of the three interglacial phases with the Eemian and the two older, Holstein 1 and Holstein 2, with the Holsteinian sensu strictu and the Do¨mnitz/ Wacken/Hoogeveen (Menke, 1968; Erd, 1970; Zagwijn, 1973) (Fig. 28.1), Preusser et al. (2005) favour a correlation of the entire Meikirch complex with MIS 7 mainly based on luminescence dating and comparison with marine climate records. In summary,

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the pollen record reveals three more or less complete interglacial phases of different climatic character, Meikirch 1, Meikirch 2 and Meikirch 3, which are separated from one another by the Birchli stadial, situated between Meikirch 1 and Meikirch 2, the latter followed by the long period of the Chutzen stadial. The Chutzen stadial is characterised by several colder and milder periods, Chutzen 1 by an open tundra, Gra¨chiswil 1 interstadial by a spread of Picea and Pinus, Chutzen 2 and Chutzen 3, the coldest phases, interrupted by the Gra¨chwil 2 interstadial which is also characterised by increases of Picea, Pinus and Betula. The Chutzen stadial is followed by the Bu¨tschwil interstadial which is characterised by an open Betula forest and remarkable amounts of Larix. It is considered to be part of the initial phase of reforestation of the Meikirch 3. Owing to the abrupt change in pollen composition, a hiatus is assumed to occur between the Bu¨tschwil interstadial and the Meikirch 3 interglacial. Concerning the interglacial sequences, Meikirch 2 is described as a temperate interglacial, Meikirch 1 and Meikirch 3 as well-developed warm periods. Preusser et al. (2005) state that Meikirch 2 might not be complete, and the climatic gradient between the Swiss alpine foreland and the Massif Central during the interglacials is not known. The sequence shows several similarities as well to the Jagonas interstadials, as well as to the Ussel and the Amargiers interstadial (Reille et al., 2000) (Table 28.2). Owing to the vegetation pattern of Meikirch 1, characterised amongst others by the lack of Fagus, a correlation with Holsteinian sequences is most unlikely. Taking all these observations as well as the unknown, perhaps even more pronounced, climatic gradient between the Alpine foreland and western central Germany into account, the Meikirch and the Scho¨ningen pollen records are tentatively compared (Table 28.3). It is not known whether the almost total lack of Abies in Meikirch 2 is due to a possible hiatus or whether it

reveals the degree of vegetation development. Its temperate character compared to that of the Meikirch 1 and 3 warm interglacials and its similarities with Middle Pleistocene interstadials of the French Central Massif, as stated above, suggests correlation with Reinsdorf Interstadials A and B. With this comparison, Meikirch 1 interglacial might correspond to the Reinsdorf interglacial. Part of the Chutzen stadial sequence might be correlated with Reinsdorf Stadial C and the cryoturbation horizon of the top sediments (Figs. 28.12, 28.13, Table 28.3). In Scho¨ningen, a peat of 1.5 m thickness topped by limnic sediments contains a pronounced interstadial, named locally the Harbke interstadial (Urban, 1996a). It has been found in a comparable lithostratigraphic position to the peat layer of the Scho¨ningen interglacial in the northern mining field Esbeck, but could not be correlated to date. The Harbke interstadial shows the following characteristics: a Betula–Pinus zone, followed by a Pinus–Betula–Picea zone and a Pinus–NAP zone with increasing amounts of Chenopodiaceae, Polygonaceae, Caryophyllaceae and Asteraceae. During pollen zone Esd V, the Pinus–Betula–Alnus– (Ericaceae–Sphagnum) zone, the growth of a local fen peat reached its climax. The Pinus–Betula–NBP zone marks the end of that biostratigraphic unit. Pollen of Larix occurs more or less continuously though with low values throughout the entire profile. If compared with the Meikirch record, the possibly truncated Bu¨tschwil interstadial might comprise part of the Harbke interstadial (Tables 28.2, 28.3). 28.3.4. Eemian (MIS 5e), Early Weichselian (MIS d, c), Late Glacial and Holocene In Scho¨ningen, the MIS 5e (Bassinot et al., 1994) equivalent travertine sediments were deposited during a period of about 6000 years, deduced from the reconstructed pollen zones (Mu¨ller, 1974a) which can be


439

Table 28.3 Tentative correlation of the Scho¨ningen Middle Pleistocene sequence with the Meikirch record (Preusser et al., 2005) Germany NE Niedersachsen

Swiss alpine foreland

Urban et al. 1988, 1991a, 1991b; Heijnis, 1992; Heijnis and Urban, 1995; Urban, 1995a, 1999, 2002

Preusser et al., 2005

Elm C Büddenstedt II Elm B Büddenstedt I Elm A stadial Schöningen 230Th/234U 180 and 227 kyr ?

Hubel Meikirch 3 Hiatus ?

Harbke interstadial

Bütschwil

Reinsdorf stadial C

(II-4/5) Chutzen sequence

Reinsdorf interstadial B

(II-4)

Reinsdorf stadial B

(II-4)

Reinsdorf interstadial A

(II-3)

kirch 2

Reinsdorf stadial A

(II-2)

Birchli Meikirch 1

Reinsdorf

230Th/234U

prox. 320 kyr?

correlated with the Eemian and Early Weichselian site of Gro¨bern (Litt, 1990) and other Northern German Late Pleistocene sequences (Behre, 1974, 1989; Menke and Tynni, 1984; Behre and Lade 1986; Urban in Veil et al., 1992; Hahne et al., 1994; Caspers, 1997). In Scho¨ningen, the terminal part of the interglacial is represented by the Carpinus–Picea phase indicating rather uniform ‘oceanic’ climatic conditions and the following Pinus–Picea–Abies phase having a more boreal and suboceanic character in Northwestern and Central Europe (Aalbersberg and Litt, 1999). Hydrological and climatic development of the late last interglacial and early glacial periods have been determined by pollen analyses, plant macro remains and molluscs (Urban et al., 1991a). The Eemian peaty layers reveal a thorium/ uranium age of 132 17 kyr (Heijnis, 1992) (Fig. 28.14, Table 28.2). The youngest sedimentation cycle in Scho¨ningen contains late Weichselian and Holocene deposits and soils, which have allowed a detailed reconstruction of Late Glacial and Holocene

MIS

7?

Mei-

7/9 ?

physical environments and human impact (Figs. 28.5, 28.14) (Thieme et al., 1987, Thieme and Maier, 1995). 28.4. SUMMARY OF STRATIGRAPHIC ¨ NINGEN ASPECTS OF THE SCHO SEQUENCE In Northwest Europe, Middle Pleistocene post-Holsteinian sequences with a definite stratigraphical position below the Early Saalian till are rather rare. Besides the previously cited correlations based on biostratigraphic findings, several other authors have correlated the Scho¨ningen sequence with long Pleistocene records (e.g. Kukla and Cilek, 1996) and placed the Holsteinian into MIS 11 and the Reinsdorf interglacial into MIS 9. Jo¨ris and Baales (2003) attempted to correlate the Scho¨ningen sequence with the marine isotope chronology using the Vostok record (Petit et al., 1999) and came to the same conclusion. In their discussion of

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the stratigraphic position and the age of the throwing spears (Thieme, 1999), which the authors wrongly placed into the Reinsdorf interglacial instead of the following interstadial (Reinsdorf interstadial B of Cycle II-4; Urban, this volume; Thieme, 1999), the authors correlate the Scho¨ningen interglacial with MIS 9a, the early Saalian Drenthe with MIS 8 and the Warthe loess with stage 6. Both the 230Th/234U ages of 180 and 227 kyr for the Scho¨ningen interglacial and of 132 17 kyr for the Eemian interglacial peat (Heijnis, 1992; Urban, 1995a) suggest a correlation of the Scho¨ningen interglacial with MIS 7 rather than with MIS 9 and the Drenthe till with MIS 6. Recently, the Antarctic Vostok ice core (EPICA Community Members, 2004; McManus, 2004) provided detailed evidence of climate development over the past 420 000 years. Of particular interest is the ascertainment that the interglacial stage MIS 11, following Termination V, was about 28 000 years long. According to Meyer (1974), Mu¨ller (1974b) and Geyh and Mu¨ller (2005), the Holsteinian interglacial as determined by counting of annual diatom layers should have covered a period of about 16 000 years. In comparison, the Rhumian interglacial was determined to have been about 25 000 years long (Mu¨ller, 1992). Those observations of the particular length of the interglacial periods and botanical pattern are used as an important argument for correlation in addition to thorium/uranium dating of Holsteinian-type locality deposits by Geyh and Mu¨ller (2005). Recent approaches of direct correlation of land–sea records between terrestrial and marine climatic indicators and ice volume proxies from deep-sea core MD01 2447 (off northwestern Iberia) show that the warmest period of MIS 11 lasted about 32 000 years (426–394 kyr) and was followed by three warm/cold cycles (394–362 kyr) (Desprat et al., 2005). During those interstadial/stadial periods, deciduous forests prevailed, and heathland in transition to open grassland characterised the cold steppe-like

stadial environments. Of utmost interest in comparison with terrestrial records of early post-Holsteinian age is pollen zone MD47S2 of the second stadial between 384 and 382 kyr, which is abruptly intercalated between interstadials MD47-I2 and MD47I3. The same is observed in NW Europe, for example, at Scho¨ningen for the postHolsteinian twofold Mißaue interstadials I and II, which are equivalent to the Dockenhuden interstadial (Hallik and Linke, 1986; Urban 1996a) and to the Pritzwalk interstadial A/B (Erd, 1973; Erd et al., 1987). They all have a sudden and short cold spell in common, dividing the interstadial into two major parts. The long warm phase of MIS 11 in northwestern Iberia is named Vigo interglacial. Desprat et al. (2005), furthermore, found that the Vigo interglacial of MIS 11 in the marine pollen record off northwestern Iberia shows a floral succession and development similar to that of the Praclaux interglacial (Reille et al., 2000) and with certain features defining the Holsteinian interglacial of Western and Central Europe. Based on those observations, the authors correlate MIS 11 with the Praclaux interglacial and the Holsteinian. The summary of Fig. 28.14 presents a subdivision of the Quaternary of Western Europe based on the biostratigraphic units of the Scho¨ningen sequence. As the 230 Th/234U age determinations on peat, which were the pioneer research of Henk Heijnis (Groningen and Sydney) in the early 1990s, are still preliminary, and as different types of sediments and soils are currently the subjects of dating processes, I propose the scheme (Fig. 28.14) based on biostratigraphic correlatives. As the efforts of determining the exact age of the Holsteinian from the marine sediments of Bossel/Germany (Geyh and Mu¨ller, 2005) suggest a correlation with MIS 9, research is now focussing on testing this age determination at different sites and on further age determination of younger interglacial and interstadial peat deposits


of definite post-Holsteinian sequences (Frechen et al., this volume). Amongst these, Scho¨ningen is of significance as it contains interglacial and interstadial peat deposits with an undisputable stratigraphical position below early Saalian till. ACKNOWLEDGEMENTS I am very indebted to Dr. Hartmut Thieme, Hannover, the archaeological excavator of the Scho¨ningen sites, for his sampling and provision of sediments as well as for fruitful and stimulating discussions, advice and financial support. I thank Christiane Hilmer, Suderburg, for valuable help with laboratory treatment of the samples and soil analyses and Barbara Albrecht for palynological work. I am very thankful to Katrin Becker and Mario Tucci, Suderburg, who helped draft the graphs and figures. I am very indebted to Helmut Mu¨ller, who gave me the opportunity to see the diagrams of pit Nachtigall and for the benefit of his comments and sharing of knowledge. Special thanks are given to Peter Kershaw, Clayton, Victoria, for critically reading the manuscript. I finally like to thank the two reviewers for their valuable advice.

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Behre, K.E., Lade, U., 1986. Eine Folge von Eem und 4 Weichsel-Interstadialen in Oerel/Niedersachsen und ihr Vegetationsablauf. Eiszeitalter und Gegenwart 36, 11–36. Bittmann, F., 1992. The Ka¨rlich Interglacial, Middle Rhine Region, Germany; vegetation history and stratigraphic position. Vegetation History and Archaeobotany 1, 243–258, Berlin. Bittmann, F., Mu¨ller, H., 1996. The Ka¨rlich Interglaical site and its correlation with the Bilshausen sequence. In: Turner, C. (Ed.), The Early Middle Pleistocene in Europe. Balkema, Rotterdam, pp. 187–193. Bo¨hme, G., 2000. Reste von Fischen, Amphibien und Reptilien aus der Fundstelle Scho¨ningen 12 bei Helmstedt (Niedersachsen) Erste Ergebnisse. Praehistoria Thuringica 4, 18–27. Boenigk, W., Frechen, M., 1998. Zur Geologie der Deckschichten von Ka¨rlich/Mittelrhein. Eiszeitalter und Gegenwart 48, 38–49. Caspers, G., 1997. Die eem- und weichselzeitliche Hohlform von Groß Todtshorn (Kr. Harburg; Niedersachsen) – Geologische und palynologische Untersuchungen zu Vegetation und Klimaverlauf der letzten Kaltzeit. In: Freund, H., Caspers, G. (Eds.), Vegetation und Palaöklima der WeichselKaltzeit im no¨rdlichen Mitteleuropa, Hannover; Schriftenreihe der deutschen Geologischen Gesellschaft 4, 7–59. de Beaulieu, J.-L., Andrieu-Ponel, V., Reille, M., Gru¨ger, E., Tzedakis, C., Svoboda, H., 2001. An attempt at correlation between the Velay pollen sequence and the Middle Pleistocene stratigraphy from central Europe. Quaternary Science Reviews 20, 1593–1602. Desprat, S., Sańchez Gon˜i, M.F., Turon, J.-L., McManus, J.F., Loutre, M.F., Duprat, J., Malaize´, B., Peyron, O., Peypouquet, J.-P., 2005. Is vegetation responsible for glacial inception during periods of muted insolation changes. Quaternary Science Reviews 24, 1361–1374. Elsner, H., 1987. Das Quarta¨r im Tagebau Scho¨ningen der Braunschweigischen Kohlen-Bergwerke AG, Helmstedt. Diplomarbeit am Fachbereich Erdwissenschaften der Universita¨t Hannover. 126 p. unpublished. Elsner, H., 2003. Verbreitung und Ausbildung Elsterzeitlicher Ablagerungen zwischen Elm und Flechtinger Ho¨henzug. Eiszeitalter und Gegenwart 52, 91–116. EPICA Community Members, 2004. Eight glacial cycles from an Antarctic ice core. Nature 429, 623–628. Erd, K., 1970. Pollen-analytical classification of the Middle Pleistocene in the German Democratic Republic, Palaeogeography, Palaeoclimatology, Palaeoecology 8, 119–132. Erd, K., 1973. Vegetationsentwicklung und Biostratigraphie der Do¨mnitz-Warmzeit (Fuhne/Saale) im

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Profil von Pritzwalk/Prignitz, Abhandlungen des Zentralen Geologischen Instituts 18, 9–48. ¨ cker-Warmzeit von Ro¨persdorf Erd, K., 1987. Die U bei Prenzlau als neuer Interglazialtyp im SaaleKomplex der DDR. Zeitschrift fu¨r Geologische Wissenschaften 15, 297–313. Erd, K., Palme, H., Pra¨ger, F.,1987. Holsteininterglaziale Ablagerungen von Rossendorf bei Dresden. Zeitschrift fu¨r Geologische Wissenschaften 15, 281–295. Felix-Hennigsen, P., Urban, B., 1982. Paleoclimatic interpretation of a thick Intra-Saalian paleosol the ‘‘bleached loam’’ on the Drenthe moraines of Northern Germany. CATENA. 9, 1–8. Field, M.H., de Beaulieu, J.-L., Guiot, J., Ponel, P., 2000. Middle Pleistocene deposits at La Coˆte, Val-de-Lans, Ise´re department, France: plant microfossil, palynological and fossil insect investigations. Palaeogeography, Palaeoclimatology, Palaeoecology 159, 53–83. Frechen, M., Sierralta, M., Oezen, D., Urban, B., 2005. Uranium series dating of peat from Central and Northern Europe. In: The Climate of Past Interglacials. Frank Sirocko, Thomas Litt, Martin Claussen (Eds.) (this volume). Geyh, M.A., Mu¨ller, H., 2005. Numerical 230Th/U dating and a palynological review of the Holsteinian/Hoxnian interglacial. Quaternary Science Reviews 24, 1861–1872. Gru¨ger, E., 1983. Untersuchungen zur Gliederung und Vegetationsgeschichte des Mittelpleistoza¨ns am Samerberg in Oberbayern. Geologica Bavarica 84, 21–40. Gru¨ger, E., Jordan, H., Meischner, D., Schlie, P., 1994. Mittelpleistoza¨ne Warmzeiten in Go¨ttingen, Bohrungen Ottostraße und Akazienweg. Geologisches Jahrbuch A 134, 167–210. Hahne, J., Kemle, S., Merkt, J., Meyer, K.-D., 1994. Eem-, weichsel- und saalezeitliche Ablagerungen der Bohrung ‘‘Quakenbru¨ck GE2’’. In: K.-D. Meyer et al. (Eds.), Neuere Untersuchungen an Interglazialen in Niedersachsen; Geol. Jb. A 134, 9–69. Hallik, R., Linke, G., 1986. Die vegetationsgeschichtliche Entwicklung des Holstein-Interglazials nach Untersuchungen in der Region Hamburg. INQUA Subcommission on European Quaternary Stratigraphy. Abstract 8, Holstein Symposium, Hamburg. Hartmann, T., 1988. Elster- bis Saale-zeitliche Sedimente im Tagebau Scho¨ningen der Braunschweigischen Kohlen-Bergwerke AG, Helmstedt. Diplomarbeit am Fachbereich Erdwissenschaften der Universita¨t Hannover. 153 p. unpublished. Heijnis, H., 1992. Uranium/thorium dating of Late Pleistocene peat deposits in N.W. Europe. PhD thesis, Rijksuniversiteit Groningen, 149 pp. Heijnis, H., Urban, B., 1995. 230Th/234U dating of the middle and late Pleistocene organic deposits from the Scho¨ningen/Helmstedt area, Lower Saxony,

Germany. Schriften der Alfred Wegener Stiftung, 2/95, 109, INQUA, XIV Congress, Berlin. Hermsdorf, N., Strahl, J., 2005. Zum Problem der sogenannten ‘‘Ueckerwarmzeit’’ (Intrasaale) Untersuchungen an neuen Bohrkernen aus dem Raum Prenzlau. Tagung der Norddeutschen Geologen, 17.-20.5.2005, Lu¨beck. Abstract, 33–34. INQUA SEQS: Subcommission on European Quaternary Stratigraphy. The Saalian sequence in the type region (Central Germany) (Halle 1992). Convention: End of Holsteinian (top) is defined as base ¨ bereinkunft: of the Saalian Complex s.l. U Obergrenze des Holstein-Interglazials entspricht Untergrenze des Saale-Komplexes. Jancyk-Kopikowa, Z., 1975. Flora interglacjalu Marzowieckiego w Fernandowie. Biuletyn Institut Geologiczny 290, 1–94. Jancyk-Kopikowa, Z., Zarski, M, 1995. The Ferdinando´w interglacial at Stanislawice near Kozienice (Central Poland). Acta Palaeobotanica, 35, 7–13. Jechorek, H., 2000. Die fossile Flora des ReinsdorfInterglazials. Palaökarpologische Untersuchungen an mittelpleistoza¨nen Ablagerungen im Braunkohlentagebau Scho¨ningen. Praehistoria Thuringica 4, 7–17. Jo¨ris, O., Baales, M., 2003. Zur Altersstellung der Scho¨ninger Speere. Vero¨ffentlichungen des Landesamtes fu¨r Archaölogie 57, 281–287. Kukla, G., Cı´lek, V. 1996. Plio-Pleistocene megacycles:record of climate and tectonics, Palaeogeography, Palaeoclimatology, Palaeoecology 35, 121–144. Lenhart, R., 1989. Schichtlagerung und Zusammensetzung Elster- bis Saale-zeitlicher Sedimente im Baufeld Esbeck, Tagebau Scho¨ningen der Braunschweigischen Kohlen- Bergwerke AG, Helmstedt. Diplomarbeit am Fachbereich Erdwissenschaften der Universita¨t Hannover. 125 p. unpublished. Lindner, L., Marciniak, B., 1998. The occurrence of four interglacials younger than the Sanian 2 (Elsterian 2) Glaciation in the Pleistocene of Europe. Acta Geologica Polonica 48, 247–263. Linke, G., Hallik, R., 1993. Die pollenanalytischen Ergebnisse der Bohrungen Hamburg-Dockenhuden (qho4), Wedel (qho2) und Hamburg Billbrook. Geologisches Jahrbuch A 138, S. 169–184. Litt, T., 1990. Pollenanalytische Untersuchungen zur Vegetations- und Klimaentwicklung wa¨hrend des Jungpleistoza¨ns in den Becken von Gro¨bern und Grabschu¨tz. Altenburger naturwissenschaftliche Forschungen 5, 92–105. McManus, J.F., 2004. A great grand-daddy of ice cores. Nature 429, 611–612. Mania, D., 1995. Die geologischen Verha¨ltnisse im Gebiet von Scho¨ningen. In: Thieme, H., Maier, R., (Eds.), Archaölogische Ausgrabungen im Braunkohlentagebau Scho¨ningen, 33–43, Hahnsche Buchhandlung, Hannover.

Interglacial Pollen Records from Scho¨ningen Mania, D., 1998. Zum Ablauf der Klimazyklen seit der Elstervereisung im Elbe-Saalegebiet. Praehistoria Thuringica 2, 5–21. Mania, D., Mai, D.-H., 2001. Molluskenfaunen und Floren im Elbe-Saalegebiet wa¨hrend des mittleren Eiszeitalters. Praehistoria Thuringica 6/7, 46–92. Menke, B., 1968. Beitra¨ge zur Biostratigraphie des Mittelpleistoza¨ns in Norddeutschland. Meyniana 18, 35–42. Menke, B., 1980. Wacken, Elster-Glazial, marines Holstein-Interglazial und Wacken-Warmzeit. In: H.E. Stremme, B. Menke (Eds.), Quarta¨r-Exkursionen in Schleswig-Holstein, Geologisches Landesamt Schleswig-Holstein. Menke, B., Tynni, R., 1984. Das Eeminterglazial und das Weichselfru¨hglazial von Rederstall/ Dithmarschen und ihre Bedeutung fu¨r die mitteleuropaïsche Jungpleistoza¨ngliederung. Geologisches Jahrbuch A 76, 120p. Meyer, K.-J., 1974. Pollenanalytische Untersuchungen und Jahresschichtenza¨hlungen an der holsteinzeitlichen Kieselgur von Hetendorf, Geologisches Jahrbuch A 21, 87–105. Mu¨ller, H., 1965. Eine pollenanalytische Neubearbeitung des Interglazial-Profils von Bilzhausen (UnterEichsfeld). Geologisches Jahrbuch 83, 327–352. Mu¨ller, H., 1974a. Pollenanalytische Untersuchungen und Jahresschichtenza¨hlungen an der eem-zeitlichenKieselgur von Bispingen/Luhe, Geologisches Jahrbuch A 21, 149–169. Mu¨ller, H., 1974b. Pollenanalytische Untersuchungen und Jahresschichtenza¨hlungen an der holstein-zeitlichen Kieselgur von Munster-Breloh, Geologisches Jahrbuch A 21,107–140. Mu¨ller, H., 1992. Climate changes during and at the end of the interglacials of the Cromerian complex, In Kukla, G.J., Wendt, E. (Eds.), Start of a Glacial, NATO ASI ser., 13, 51–69. Mu¨ller, H., Ho¨fle, H.-C., 1994. Das HolsteinInterglazialvorkommenbei Bossel westlich von Stade und Wanho¨den no¨rdlich Bremerhaven, Geologisches Jahrbuch, A 134, 71–116. Petit J.R., Jouzel, J., Barkov, N.I., Barnola, J.-M., Basile, I., Bender, M., Chappellaz, J., Davis, M., Delaygue, G., Delmotte, M., Kotlyakov, V.M., Legrand, M., Lipenkov, V.Y., Lorius, C., Pepin, L., Ritz, C., Saltzman, E, Stievenard, M., 1999. Climate and atmospheric history of the past 420000 years from the Vostok ice core, Antarctica. Nature 399, 429–436. Preusser, F., Drescher-Schneider, R., Fiebig, M., Schlu¨chter, C., 2005. Re-interpretation of the Meikirch pollen record, Swiss Alpine Foreland, and implications for Middle Pleistocene chronostratigraphy. Journal of Quaternary Science 20, 607–620. Reille, M., de Beaulieu, J.-L., 1995. Long Pleistocene Pollen Records from the Praclaux Crater, SouthCentral France. Quaternary Research 44, 205–215.

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Reille, M., Andrieu, V., de Beaulieu, J-L., Guenet, P., Goeury, C., 1998. A long pollen record from Lac Du Bouchet, massif Central, France: For the Period ca. 325 to 100BP (OIS 9 c to OIS 5e), Quaternary Science Reviews 17, 1107–1123. Reille, M., de Beaulieu, J.-L., Svoboda, V., AndrieuPonel, V., Goeury, C., 2000. Pollen analytical biostratigraphy of the last five climatic cycles from a long continental sequence from the Velay region (Massif Central, France). Journal of Quaternary Science 15, 665–685. Richter, D., 1998. Thermolumineszenzdatierungen erhitzter Silices aus mittel- und jungpalaölithischen Fundstellen. Anwendung und methodische Untersuchungen. PhD Thesis, Universita¨t Tu¨bingen. Sarnthein, M., Stremme, H.-E., Mangini, A. 1986. The Holsteinian Interglacial: time- stratigraphic position and correlation to stable-isotope stratigraphy of deep-sea sediments. Quaternary Research 26: 283–298. Schoch, W.H., 1999. Holz als Informationstra¨ger aus dem Palaölithikum. Praehistoria Thuringica 3, 98–106. Strahl, J., 1999. Biostratigraphische Untersuchungen im Bereich des Oberen Allertales (Raum Morsleben und Ummendorf). 66, Tagung AG Nordwestdt. Geologen, Tagungsband und Exkursionsfu¨hrer, 119–124. Halle. Thieme, H., 1996. Altpalaölithische Wurfspeere aus Scho¨ningen, Niedersachsen – Ein Vorbericht, Archaölogisches Korrespondezblatt 26, 377–393. Thieme, H., 1997. Lower Paleolithic hunting spears from Germany. Nature 385, 807–810. Thieme, H., 1998. Altpalaölithische Wurfspeere von Scho¨ningen, Niedersachsen. Praehistoria Thuringica 2, 22–31. Thieme, H., 1999. Altpalaölithische Holzgera¨te aus Scho¨ningen, Lkrs. Helmstedt. Bedeutsame Funde zur Kulturentwicklung des fru¨hen Menschen. Germania 77, 451–487. Thieme, H., Mania, D., 1993. ‘‘Scho¨ningen 12’’ – ein mittelpleistoza¨nes Interglazialvorkommen im Nordharzvorland mit palaölithischen Funden, Ethnographisch-Archaölogische Zeitschrift 1993, 34: 610–619 Thieme, H., Maier, R., 1995. Archaölogische Ausgrabungen im Braunkohlentagebau Scho¨ningen, Landkreis Helmstedt. 191 pp. Braunschweigische Kohlen-Bergwerke AG Helmstedt, (Eds.), Verlag Hahnsche Buchhandlung Hannover. Thieme, H., Maier, R., Urban, B., 1987. Archaölogische Schwerpunktuntersuchungen im Helmstedter Braunkohlenrevier (ASHB) – zum Stand der Arbeiten 1983–1986. Archaölogisches Korrespondenzblatt 17, 445–462. Thieme, H., Maier, R., Urban, B., 1992. Neue Erkenntnisse zum urgeschichtlichen Siedlungsgeschehen. – Archaölogie in Deutschland, Heft 2, 26–30.

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Thieme, H., Mania, D., 1993. ‘‘Scho¨ningen 12’’ – ein mittelpleistoza¨nes Interglazialvorkommen im Nordharzvorland mit palaölithischen Funden, Ethnographisch-Archaölogische Zeitschrift 34, 610–619. Thieme, H., Mania, D., Urban, B., van Kolfschoten, T., 1993. Scho¨ningen (Nordharzvorland) eine altpalaölithische Fundstelle aus dem mittleren Eiszeitalter, Archaölogisches Korrespondenzblatt 23, 147–163. Tschee, W., 1991. Die pleistoza¨ne Schichtfolge im Tagebau Scho¨ningen Baufeld Esbeck der Braunschweigischen Kohlen-Bergwerke AG, Helmstedt. Diplomarbeit am Fachbereich Erdwissenschaften der Universita¨t Hannover. 75p. unpublished. Turner, C., 1970. The Middle Pleistocene deposits of Marks Tey, Essex. Philosophical Transactions of the Royal Society of London, Series B 257, 3373–440. Turner, C., 1998. Volcanic maars, long Quaternary sequences and the work of the INQUA Subcommission on European Quaternary stratigraphy. In: Cavarretta, G., Fornaceri, M., Follieri, M., Girotti, Turner, C. (Guest Eds.), Quaternary Stratigraphy in Volcanic Areas. Quaternary International 47/48, 3–20. Urban, B. (1983). Biostratigraphic correlation of the Ka¨rlich Interglacial, Northwestern Germany. BOREAS, 12, pp. 83–90, Oslo. Urban, B., 1992. Interglacial/glacial transitions recorded from middle and young Pleistocene sections of eastern Lower Saxony-Germany. In: Kukla, G.J., Went, E. (Eds.), Start of a Glacial. NATO ASI Series, Vol. I 3, 37–50, Springer Verlag, Berlin. Urban, B., 1995a. Palynological evidence of younger Middle Pleistocene Interglacials (Holsteinian, Reinsdorf, Scho¨ningen) in the Scho¨ningen open cast lignite mine (eastern Lower Saxony/ Germany).Mededelingen Rijks Geologische Dienst 52, 175–186. Urban, B., 1995b. Vegetations- und Klimaentwicklung des Quarta¨rs im Tagebau Scho¨ningen. In: Thieme, H., Maier, R., (Eds.), Archaölogische Ausgrabungen im Braunkohlentagebau Scho¨ningen, pp. 44–56, Hahnsche Buchhandlung, Hannover. Urban, B., 1996a. Mittelpleistoza¨ne Waldzeiten im Tagebau Scho¨ningen: Spektren aus dem HolsteinInterglazial und dem Harbke-Interstadial. In: Spuren der Jagd – Die Jagd nach Spuren. Tu¨binger Monographien zur Urgeschichte 11, 487–495. Urban, B., 1996b. Zur Palaöo¨kologie und Stratigraphie des Mittelpleistoza¨ns im Tagebau Scho¨ningen/NO Niedersachsen. Landesamt fu¨r Natur und Umwelt des Landes Schleswig-Holstein: Bo¨den als Zeugen der Landschaftsentwicklung, 127–140. Urban, B., 1997. Grundzu¨ge der eiszeitlichen Klimaund Vegetationsgeschichte in Mitteleuropa. In: Wagner, G.A., Beinhauer, K.W. (Eds.), Homo

heidelbergensis von Mauer – Das Auftreten des Menschen in Mitteleuropa, 241–265, Universita¨tsverlag C. Winter, Heidelberg. Urban, B., 1999. Middle and Late Pleistocene biostratigraphy and paleoclimate of an open-pit coal mine Schoningen: Germany. Chinese Science Bulletin 44 Suppl., 30–37. Urban, B., 2002. Rekonstruktion pleistoza¨ner und holoza¨ner Landschafts- und Klimageschichte im no¨rdlichen Mitteleuropa mit Hilfe limnisch-telmatischer und terrestrischer Sediment- und Bodenabfolgen. In: Geo 2002 – Planet Erde: Vergangenheit, Entwicklung, Zukunft. Deutsche Geologische Gesellschaft 21, 336–337. Urban, B., Thieme, H., Elsner, H., 1988. Biostratigraphische, quarta¨rgeologische und urgeschichtliche Befunde aus dem Tagebau ‘‘Scho¨ningen’’, Landkrs. Helmstedt. Zeitschrift der deutschen geologischen Gesellschaft 139, 123–154. Urban, B., Elsner, H., Ho¨lzer, A., Mania, D., Albrecht, B., 1991a. Eine eem- und fru¨hweichselzeitliche Abfolge im Tagebau Scho¨ningen, Landkreis Helmstedt. Eiszeitalter und Gegenwart 41, 85–99. Urban, B., Lenhard, R., Mania, D., Albrecht, B., 1991b. Mittelpleistoza¨n im Tagebau Scho¨ningen, Ldkrs. Helmstedt. Zeitschrift der deutschen geologischen Gesellschaft 142, 351–372. van den Boogard, C., Boogard, P., v.d., Schmincke, H.-U.,1989. Quarta¨rgeologisch-tephrostratigraphische Neuaufnahme und Interpretation des Pleistoza¨nprofils Ka¨rlich. Eiszeitalter und Gegenwart 39, 62–86. van Kolfschoten, T., 1995. Faunenreste des altpalaölithischen Fundplatzes Scho¨ningen 12 (Reinsdorf Interglazial). In: Thieme, H., Maier, R., (Eds.), Archaölogische Ausgrabungen im Braunkohlentagebau Scho¨ningen, Hahnsche Buchhandlung, Hannover, 85–94 Veil, S.; Breest, K.; Ho¨fle, H.-C.; Meyer, H.-H.; Plisson, H.; Urban-Ku¨ttel, B.; Wagner, G.A.; Zo¨ller, L., 1992. Ein mittelpalaölithischer Fundplatz aus der Weichsel-Kaltzeit bei Lichtenberg, Lkr. Lu¨chow-Dannenberg. Germania 72, 1–66. Welten, M., 1982. Pollenanalytische Untersuchungen im ju¨ngeren Quarta¨r des no¨rdlichen Alpenvorlandes der Schweiz. Beitra¨ge zur Geologischen Karte der Schweiz – Neue Folge 156, 179pp. Welten, M., 1988. Neue pollenanalytische Ergebnisse u¨ber das ju¨ngere Quarta¨r des no¨rdlichen Alpenvorlandes der Schweiz (Mittel- und Jungpleistoza¨n). Beitra¨ge zur Geologischen Karte der Schweiz – Neue Folge 162, 40 pp. Zagwijn, W.H., 1973. Pollenanalytic studies of Holsteinian and Saalian beds in the northern Netherlands. Mededelingen Rijks Geologische Dienst 24, 139–156.

29. Mammalian Faunas From the Interglacial Periods in Central Europe and Their Stratigraphic Correlation Wighart von Koenigswald Institut fu¨r Palaöntologie der Universita¨t Bonn, Nussallee 8, D-53115 Bonn

ABSTRACT During the Pleistocene, climatic oscillations caused multiple faunal exchanges in Central Europe. The interglacial Elephas antiquus assemblage reinvaded Central Europe several times. The various Middle Pleistocene interglacial periods are very difficult to distinguish based on the fauna alone. The classical biostratigraphic tool, the different levels in a continuous evolutionary sequence, can only be applied in very few genera, e.g. Arvicola, to a limited degree. During the Middle and Late Pleistocene, a decline in faunal diversity can be observed. 29.1 INTRODUCTION The mammalian fauna of the Central European Pleistocene is characterized by drastic turnovers induced by climatic changes. During the Middle and Late Pleistocene, a continental fauna characterized by Mammuthus and Coelodonta characterizes the cold periods. During the warm periods, the interglacials, a fauna dominated by Elephas antiquus was present. In the Middle and Late Pleistocene, a Mammuthus assemblage can clearly be differentiated from an Elephas assemblage. Only very few herbivores occur in both assemblages (Fig. 29.1). With each climatic change, the fauna was exchanged. That means that the new fauna invaded, while the previous one became locally extinct (von Koenigswald, 2003). This faunal exchange of the Middle and Late Pleistocene is more significant than in most other areas of the world, due to the specific geographical situation of Central (and Western)

Europe. The West–East oriented mountain ranges of the Pyrenees and Alps stabilized the climate in the Mediterranean as well as in Central Europe. During cold periods, small oscillations of warmer climate were buffered, and thus the glacial fauna was not changed during most interstadials. Only interglacials led to an intensive faunal exchange, and the Mediterranean fauna expanded its area using the passages of the Rhone and the Danube valleys (Fig.29.2). Temperature and humidity are the main components of climate. The very significant difference between glacial and interglacial faunas in Central Europe is due to the shift from a highly continental biome during glacial periods to a more maritime-influenced climate during the interglacials. Thus, it is the changing precipitation, which has the greatest influence. Faunal lists for most of the sites mentioned here were given in von Koenigswald and Heinrich (1999), including the references to these local faunas, and thus have not been repeated here. 29.2 THE EARLY PLEISTOCENE The climatic fluctuations characterizing the Pleistocene had already started in the Pliocene, as documented in the pollen record. In the mammalian record, no typical cold faunas are yet known in the fossil record from the Pliocene or Early Pleistocene. This might be due to the very scarce fossil record, or to the fact that the mammalian fauna was not yet adapted to a cold environment. At least, there is no indication of faunal exchanges as in the Middle and Late Pleistocene.

446

Wighart von Koenigswald

Elephas - fauna

Mammuthus - fauna

Crocuta c. spelaea Ovibos moschatus

Dama dama Capreolus capreolus

Panthera leo spelaea

Bubalus murrenis

Rangifer tarandus

Coelodonta antiquitatis Bos primigenius

Ursus spelaeus

Hippopotamus amphibius Bison priscus

Saiga tatarica

Stephanorhinus kirchbergensis Cervus elaphus

Sus scrofa Elephas antiquus

Equus ferus

Megaloceros giganteus

Mammuthus primigenius

Fig. 29.1 The Elephas assemblage and the Mammuthus assemblage occurred alternatively in Central Europe during the Middle and Late Pleistocene. Only few herbivores but several carnivores occurred during interglacial and glacial environments (von Koenigswald, 2002).

The mammalian fauna of Untermaßfeld near Meiningen represents temperate climatic conditions of a very late phase of the Early Pleistocene. The rich fauna, containing many large herbivores and magnificent carnivores, was deposited in fluvial sediments after a great flood (Kahlke, 1997–2001) and includes Hippopotamus amphibius, a most indicative taxon for interglacial conditions. The specific ecological requirements of Hippopotamus will be discussed in the section on the Eemian. 29.3 MIDDLE PLEISTOCENE INTERGLACIAL FAUNAS BEFORE THE ELSTERIAN Among the faunas of the early Middle Pleistocene, two stratigraphic levels can be distinguished according to the occurrence of

two large vole taxa (Fig. 29.3). The genus Mimomys has rooted molars and characterizes the older faunas, while the more derived Arvicola characterized by rootless molars is present in the younger part of the Middle Pleistocene, the Late Pleistocene and Holocene. The genus Arvicola shows some progressive evolution during that time interval and thus is an important stratigraphic tool. The Mimomys savini faunas are well documented in Thu¨ringen, especially at Voigtstedt (Kahlke, 1965) and Su¨ssenborn (Kahlke, 1969). They are younger than the Matuyama/Brunhes boundary and thus belong to the early Middle Pleistocene. The fauna of Voigtstedt represents most probably a temperate environment, as indicated by the presence of Sus scrofa, Capreolus suessenbornensis and a flying squirrel Petauria voigtstedtensis. Voigtstedt is correlated with

Mammalian Faunas From the Interglacial Periods

?

Area of temporal occurence

Core

447

Core area of the glacial fauna

?

? area o f

the interg

lacial fauna 500 km

?

Fig. 29.2 Immigration routes of the Elephas assemblage from the South and the Mammuthus assemblage from the North East during the Middle and Late Pleistocene.

Geology

Small mammal stratigraphy

Holocene

Eemian

Arvicola cantianus-terrestris faunas

Elephas antiquus Hippopotamus Bubalus

5e

Lehringen Taubach

Warthe Drente Dömintz

Late Arvicola cantianus faunas

Elephas antiquus Bubalus

?7

Elephas antiquus

?9 ?11

W. Ehringsdorf Steinheim/Murr Schöningen Bilzingsleben

Hosteinian Elsterian Early Arvicola cantianus faunas Cromerian Complex

Lower Pleistocene

Important faunas

1

Arvicola terrestris-faunas

Weichselian Late Pleistocene

Middle Pleistocene

Significant Interglacials immigrants (and MIS)

Brunhes Matuyama

Arvicola Mimomys +

Elephas antiquus

Kärlich G Mosbach Mauer

Elephas antiquus Hippopotamus Arvicola

Mimomys savini-faunas

Mimomys savini-faunas with M. pusillus

Süssenbron Voigtstedt Hippopotamus

>19

Untermaßfeld

Fig. 29.3 Biostratigraphy of the Middle and Late Pleistocene in Central Europe (von Koenigswald and Heinrich, 1999).

448


the West Runton Freshwater Bed of the Cromer Forestbed Formation in East Anglia (Stuart, 1981). The fauna of Su¨ssenborn contains Ovibos and Rangifer, indicators of a cold climate as in later faunas and at present. Whether these genera represent an arctic environment in the early Middle Pleistocene remains uncertain since other taxa, especially the rhinos present in this fauna, are well known from warmer periods. The interglacial faunas from Jockrim in the Pfalz and from Wu¨rzburg-Schalksberg cannot be correlated stratigraphically, since neither Mimomys nor Arvicola is present. Both sites produced Hippopotamus amphibius, a faunal element occurring only during interglacial climatic conditions. At Mauer near Heidelberg, an interglacial period is represented, which is distinctly younger than Voigtstedt and Su¨ssenborn, since Arvicola cantianus has replaced Mimomys savini. Within the Arvicola cantianus faunas, Mauer represents an early stage since the diversity of insectivores and rodents is quite high. Together with Arvicola cantianus, another important faunal element, the straight-tusked elephant Elephas antiquus, occurs here for the first time. This elephant will be significant for all later interglacial faunas in Central Europe (except the Holocene). It should be mentioned that Mauer is the type locality of Homo heidelbergensis. The fully interglacial character of this site is stressed by the presence of Hippopotamus amphibius besides Sus scrofa and Capreolus capreolus priscus. The fauna of Mauer represents one of the interglacial periods within the later part of the Cromerian Complex, but it is not necessarily the latest interglacial. Another interglacial assemblage representing the early Arvicola cantianus faunas with a great diversity of insectivores and rodents, e.g. containing Talpa minor and Pliomys, was described from Ka¨rlich G. This site is important since the various layers are intercalated with tuffs. A tuff higher up in the profile was dated at 618 ka. A similar

and very rich early Arvicola cantianus fauna was found in the site Miesenheim nearby, covered by the same tuff. Distinctly younger is the so-called ‘Ka¨rlich Interglacial’ which was detected at the site Ka¨rlich Seeufer. The lake sediments produced remains of six individuals of Elephas antiquus together with artifacts (Bosinski, 1995). Similarly to the somewhat older fauna of Ka¨rlich H, it represents the early Arvicola cantianus fauna. The underlying ‘Brocken tuff’ provided an age of 396 ka. According to the pollen assemblages, Ka¨rlich Seeufer was regarded as late Cromerian and correlated convincingly with Bilshausen (Bittmann and Mu¨ller, 1996), which definitively underlies sediments of the Elsterian. Thus, the early Arvicola cantianus faunas antedate the Elsterian and represent part of the Cromerian Complex, but this faunal type might have occurred during several interglacial periods with cold or even glacial conditions in between. Some arctic faunal elements, Ovibos and Rangifer, occur in Mosbach 2, but typical glacial faunas are missing so far. The mammalian fauna of the Elsterian is more or less unknown. The occurrence of Coelodonta antiquus and Rangifer tarandus in Bornhausen near Seesen constitutes one of the faunas correlated with this glacial period. 29.4 MIDDLE PLEISTOCENE INTERGLACIAL FAUNAS BETWEEN ELSTERIAN AND SAALIAN Biostratigraphers struggle with the number of interglacial phases after the Elsterian and before the Eemian (Fig. 29.3). Mania and Thomae (2006) postulate four independent interglacial phases within the Holstein Complex and two additional interglacial periods within the Saalian. Sarntheim et al. (1986) and Schreve (2001) correlate the Holsteinian with MIS 11 and accordingly count three interglacials for this time period. Geyh and Mu¨ller (2005) dated the type locality of the Holsteinian and correlated it with MIS 9.


Litt et al. (2005) accepted one or two interglacial phases after the Holsteinian and before the first Saalian ice advance (Drenthe), rejecting any interglacial phase between the Drenthe and the Warthe. Thus besides the Holsteinian, only one (or two) additional interglacial phases are accepted as antedating the Eemian. Various sites have produced interglacial faunas, such as Bilzingleben II, Scho¨ningen 12, Scho¨ningen 13, Steinheim/Murr and Weimar-Ehringsdorf, but none of them shows a complete interglacial cycle in their pollen record. Bilzingsleben II is definitively younger than the Elsterian, but according to Mania (1997) and Mania and Thomae (2006), it is not the first interglacial period after the Esterian. The horizon, which produced the remains of Homo ‘erectus’ bizingslebenensis, provided a typical interglacial fauna with Elephas antiquus, Stephanorhinus kirchbergensis, Sus scrofa, and Capreolus. Among the rodents, the vole Arvicola cantianus is present, but some important elements from the early Arvicola cantianus fauna, e.g. Talpa minor, Drepanosorex and Pliomys, are missing although the fauna is fairly rich. Therefore, this faunal association is regarded as the younger Arvicola cantianus fauna. Another difference is that the rhino Stephanorhinus hundsheimensis is replaced by the more derived S. kirchbergensis. The faunas of Scho¨ningen 12 and 13 fit into the same faunal pattern, but based on the faunal content it is unclear whether they represent the same interglacial period or a slightly younger one. Differences seen in horses compared to those in Bilzingleben II are very difficult to evaluate, due to the limited material and the great tendency of horses to develop local forms. It is possible that Bos primigenius is present (E.v.Asperen, personal communication) which might be important for the comparison with Steinheim. The river deposits of Steinheim/Murr provided a very rich interglacial Elephas antiquus assemblage, which has often been

449

correlated with the Holsteinian (Adam, 1954, 2003). This was based on the old scheme with only three possible interglacials. Since the true number of interglacials during the Middle Pleistocene is known from the deep-sea record, a necessary correlation with the Holsteinian is obsolete. This Steinheim interglacial is older than the Eemian, and according to the geology it might correlate with MIS 7. The skull of Homo steinheimensis differs distinctly from the human remains found in Bilzingsleben II, and thus it belongs most probably to another human immigration. In the fauna, two important bovids occur for the first time: Bos primigenius and Bubalus murrensis. They are not present in Bilzingsleben II. The water buffalo is of great ecological significance since this animal requires open water during the winter to shelter from cold winds. The specific ecological requirements of Bubalus will be discussed together with those of Hippopotamus in the section on Eemian faunas. Weimar-Ehringsdorf is a travertine deposit, predominately of interglacial character. Human remains and a rich fauna were excavated from the lower travertine, which is of interest here. Traditionally, the site was regarded as Eemian, but several arguments indicate that this fauna represents an older interglacial period. Arvicola shows an evolutionary level intermediate between the one from Bilzingsleben and those of typical Eemian sites, for example Taubach and Burgtonna. Absolute dates indicate a correlation with MIS 7 (Schreve and Bridgland 2002). The difficulty of biostratigraphical correlation is caused by the similarity of the typical Elephas antiquus assemblages from the various interglacial periods of the late Middle Pleistocene and the Eemian. Differences in evolutionary level are restricted to few taxa. The mammalian fauna indicates that Bilzingsleben is earlier than Steinheim. Scho¨ningen may be the equivalent of Steinheim. Whether Weimar-Ehringsdorf belongs to the same interglacial cannot be decided from the faunal record.

450


Although one has to assume that the Elephas antiquus assemblages immigrated repeatedly, differences in the fauna are so limited that a biostratigraphic differentiation is not yet possible. 29.5 MAMMALIAN FAUNAS FROM THE EEMIAN After the end of the Saalian, the Elephas antiquus assemblage reinvaded from the Mediterranean. The mammalian fossil record is not rich enough to demonstrate the sequence of immigration similarly to that demonstrated among various taxa of the vegetation (Litt, 1994). The Elephas antiquus assemblage of the last interglacial does not show significant evolutionary changes compared to the previous interglacials. One of the few guide fossils is the vole Arvicola, which is characterized in its evolution by the transitional stage from A. cantianus to A. terrestris. Two rhinos Stephanorhinus kirchbergensis and S. hemitoechus are present. On the continent, S. kirchbergensis dominates over the other, while S. hemitoechus is the only rhino on the British Isles during this interglacial. In Northern Germany, the Eemian provides several localities with the typical Elephas antiquus assemblage including Stephanorhinus kirchbergensis, Sus scrofa, Dama dama, Cervus elaphus; Capreolus capreolus; Megaloceros giganteus and Bos primigenius. These faunas contain Equus ferus and sometimes Equus hydruntinus as well, species which are often claimed to represent a steppic environment. But obviously they occurred in forested biomes as well, since a dense forest can be assumed from the pollen record (Litt, 2000). In Thu¨ringen, the travertines of Burgtonna (Kahlke, 1978) and Taubach (Kahlke, 1977) provided important faunal remains of the Elephas antiquus assemblage. The stratigraphic position of the lakesite Neumark-Nord is debated. It is superimposed on the moraines of the Drenthe

glaciation (lower Saalian). Mania and Thomae (2006) postulate that the lake basin of Neumark-Nord 1 represents the first intra-Saalian interglacial phase covered by Warthe deposits (late Saalian glacial expansion). But Litt (in Mania et al., 1990) places Neumark-Nord into the Eemian according to the development of the vegetation and denies an intra-Saalian interglacial after the Drenthe. In the rodent fauna, the occurrence of Apodemus maastrichtiensis has to be mentioned. It is present at Weimar-Ehringsdorf but was not found in Eemian faunas of the area (Heinrich, 2001). But the stratigraphic value of this rodent is still open. Neumark-Nord 1 produced a number of well-preserved and entire skeletons of interglacial mammals, e.g. Elephas antiquus; Stephanorhinus kirchbergensis and Bos primigenius. The interglacial Dama dama geiselana is represented by a great number of male individuals with impressive antlers (Pfeiffer, 1999). Cervus elaphus; one of the few herbivores occurring during interglacial and glacial phases, is well represented, too. It is assumed that algal blooms of cyanobacteria poisoned at least some of these animals (Braun and Pfeiffer, 1994) (Tables 29.1, 29.2). Deposits of the Rhine River in the northern Oberrheinebene near Darmstadt are intensively quarried and produced a rich glacial fauna and several interglacial species (von Koenigswald, 1988). The stratigraphy of the pits cannot be observed directly since the profile is under the water table. Nevertheless, long-time observation indicates that the glacial fauna comes from the upper part while thick black trunks of oak (Quercus sp.) characterize the lower section. The interglacial fauna is quarried from the same depth. Most probably, this fauna belongs to the last interglacial, the Eemian. Typical and frequent faunal elements are those of Elephas antiquus, Stephanorhinus kirchbergensis, Dama dama and Bos primigenius. In addition, Bubalus murrensis and Hippopotamus amphibius


451

Rheinsande

Burgtonna

Taubach

MIS 1 Holocene

Weichselian

Saalian

Eemian MIS 5e

Lehringen

WeimarEhringsdorf

Steinheim/Murr

Schöningen II

MIS 9- MIS 7

Bilzingleben II

Kärlich G

Elsterian

Middle Pleistocene Mosbach II

Mauer

Voigtstedt

Untermaßfeld

Early Pleistocene

Table 29.1 Occurrence of selected large mammals in representative interglacial sites of Germany

Elephas antiquus Stephanorhinus etruscus Stephanorhinus hundsheimensis Stephanorhinus kirchbergensis

sp.

Stephanorhinus hemitoechus m.

Equus div. sp.

m.

sp.

st.

t.

Equus hydruntinus Hippopotamus amphibius Sus scrofa Dama reichenaui

sp.

sp.

Dama dama

sp.

sp.

( )

Cervus elaphus Megaloceros verticornis Megaloceros giganteus Alces latifrons

sp.

c.

sp.

Alces alces Capreolus suessenbornensis Capreolus capreolus Bison schoetensacki Bison priscus

sp.

m.

Bison bonasus Bos primigenius

?

Bubalus murrensis

occur in several of these sand pits. Their preservation and frequency exclude a redeposition from older sediments. A correlation with the last interglacial is plausible

due to the geological situation and the rich occurrence of Hippopotamus during the timeequivalent Ipswichian in the British Isles. In the same sediments, Alces latifrons and

452


MIS 1 Holozän

Weichselian Rheinsande

Burgtonna

Taubach

Eemian MIS 5e

Lehringen

WeimarEhringsdorf

Steinheim/Murr

Schöningen II

Bilzingleben II

MIS 9- MIS 7

Saalian

Elsterian Kärlich G

Mosbach II

Middle Pleistocene Mauer

Voigtstedt

Untermaßfeld

Early Pleistocene

Table 29.2 Occurrence of selected small mammals in representative interglacial sites of Germany

Talpa minor Talpa europaea Sorex (D.) savini

sp.

Mimomys savini Arvicola cantianus Arvicola cantianus/terrestris Arvicola terrestris Pliomys div sp. Trogontherium cuvieri

?

Castor fiber Apodemus sylvaticus Apodemus maastrichensis

sp. aff.

Trogotherium cuvieri were found (von Koenigswald and Menger, 1997). These species were often regarded as typical for much older deposits. But Alces latifrons is known from the interglacial of Weimar-Ehringsdorf, and the fossil limit of Trogontherium has become younger and younger during recent decades. Thus, the use of last appearance dates for stratigraphic purposes may be questionable, especially for rare species. The occurrence of Bubalus and Hippopotamus is of great ecological significance. Their present distribution in subtropical regions does not indicate similar conditions for periods when these animals were living in Central Europe. These animals can tolerate lower temperatures but submerge in the water to escape from cold winds. During winters, open water in rivers and lakes was only available in the Rhine valley, when the maritime influence was significantly

higher than today. That means mild winters but cooler summers if the annual mean temperature was not raised very much; and from the vegetation we know that the annual temperature was only 2 or 3 higher than today. Thus, a high maritime influence on the climate can be assumed for the Rhine area at least during part of the Eemian (von Koenigswald, 1988, 1991). The maritime influence was certainly less towards the East; at least Hippopotamus was not found further East during the Eemian. In this context, it is worthwhile noticing that the easternmost occurrence of this species is near Warsaw; thus it seems to avoid a very continental climate. This example indicates that it is not only the temperature, but even more the precipitation, which controls the faunal composition of the interglacial periods. In each Middle Pleistocene interglacial period and in the Eemian, the Elephas


antiquus assemblage was present in Central Europe. Since there was no major refuge north of the Alpine arc the species had to re-immigrate each time from the Mediterranean. The composition of the Elephas antiquus assemblage is fairly similar in various interglacial periods, but some species occur irregularly, or only in specific interglacial periods. The two genera Hippopotamus and Bubalus were mentioned because of their ecological significance. Finds of Macaca sylvanus are very rare. The faunal record of Central Europe does not allow us to decide whether two, three or even four different interglacial phases were intercalated between the Elsterian and the Saalian. 29.6 THE WEICHSELIAN AND THE HOLOCENE With the beginning of the last glaciation, the interglacial fauna became locally extinct in Central Europe. Most probably, the species survived longer in the Mediterranean, but their disappearance is not well documented, neither stratigraphically nor regionally. During the Weichselian, interglacial species such as Sus scrofa may have reinvaded Central Europe only sporadically (von Koenigswald and Heinrich, 1996). While Elephas antiquus and Stephanorhinus kirchbergensis became extinct, Dama and Hippopotamus survived in the eastern Mediterranean, but these species were not able to re-colonize Central Europe during the Holocene. Dama reinvaded Central and Western Europe in each interglacial period but not during the Holocene, although the climatic conditions were more favourable during the middle Holocene (Atlanticum) than today. However, when human reintroduced Dama in medieval times, this deer flourished and is now widespread again (von Koenigswald, 2002). This example indicates that it is not only the climate, through temperature and humidity, which

453

controls the faunal composition of interglacial periods. There are particular historical factors involved too.

REFERENCES Adam K.D., 1954. Die mittel-pleistoza¨nen Faunen von Steinheim an der Murr (Wu¨rtemberg). Quaternaria 1, 131–144. Adam, K.D., 2003. Der Homo steinheimensis im Spannungsfeld von Alt- und Neumensch. Vero¨ffentlichungen des Landesamtes fu¨r Archaölogie Sachsen-Anhalt - Landesmuseum fu¨r Vorgeschichte 57. Bittmann, F., Mu¨ller, H., 1996. The Ka¨rlich Interglacial site and its correlation with the Bilshausen sequence. In: Turner, C. (Ed.), The early Middle Pleistocene in Europe. Balkema, Rotterdem, 187–193. Bosinski, G., 1995. The Palaeolithic and Mesolithic of the Rhineland. In: Schirmer, W. (Ed.), Quaternary fieldtrips in Central Europe 2 Field trips on special topics. Pfeil, Mu¨nchen, 829–999. Braun, A., Pfeiffer, T., 1994. Cyanobacterial blooms as the cause of a Pleistocene large mammal assemblage. Paleobiology 28, 139–154. Geyh, M.A., Mu¨ller, H., 2005. Numerical 239Th/U dating and a palynological review of the Holsteinian/Hoxnian interglacial. Quaternary Science Reviews 24, 1861–1872. Heinrich, W.D., 2001. Kleinsaügerreste aus interglazialen Ablagerungen von Neumark-Nord, Mitteldeutschland. Praehistoria Thuringica 6/7, 132–138. Kahlke, H.D. (Ed.), 1965. Das Pleistoza¨n von Voigtstedt. Palaöntologische Abhandlungen, Palaözoologie II, 221–692. Kahlke, H.D. (Ed.), 1969. Das Pleistoza¨n von Su¨ssenborn. Palaöntologische Abhandlungen, Palaözoologie III, 367–788. Kahlke, H.D. (Ed.), 1977. Das Pleistoza¨n von Taubach bei Weimar. Quarta¨rpalaöntologie 2, 1–509. Kahlke, H.D. (Ed.), 1978. Das Pleistoza¨n von Burgtonna. Quarta¨rpalaöntologie 3, 1–399. Kahlke, R.D. (Ed.), 1997–2001. Das Pleistoza¨n von Untermaßfeld bei Meiningen (Thu¨ringen) Teil 1– 3, Ro¨misch-Germanisches Zentralmuseum Mainz, Monographien 40/1–3. von Koenigswald, W. (Ed.), 1988. Zur Palaöklimatologie des letzten Interglazials im Nordteil der Oberrheinebene. Palaöklimaforschung 4, 1–327. von Koenigswald, W., 1991. Exoten in der GroßsaügerFauna des letzten Interglazials von Mitteleuropa. Eiszeitalter und Gegenwart 41, 70–84. ¨ kologie und Biosvon Koenigswald, W. 1992. Zur O tratigraphie der beiden pleistoza¨nen Faunen von

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Mauer bei Heidelberg. In: Beinhauer, K.W., Wagner, G.A. (Eds), Schichten von Mauer - 85 Jahre Homo erectus heidelbergensis. – Mannheim, 101–110. von Koenigswald, W. 2002. Lebendige Eiszeit Klima und Tierwelt im Wandel. Wissenschaftliche Buchgesellschaft Darmstadt und Theiss Stuttgart, pp. 190. von Koenigswald, W. 2003. Mode and cause for the Pleistocene turnovers in the mammalian fauna of Central Europe. Deinsia 10, 305–312. von Koenigswald, W., Heinrich, W.-D., 1996. Kurze Charakterisierung der Vera¨nderungen in der Saügetierfauna des Jungquarta¨rs in Mitteleuropa. Tu¨binger Mongraphien zur Urgeschichte 11, 437–448. von Koenigswald, W., Heinrich W.-D., 1999. Mittelpleistoza¨ne Saügetierfaunen aus Mitteleuropa der Versuch einer biostratigraphischen Zuordnung. Kaupia 9, 53–112. von Koenigswald, W., Menger, F., 1997. Mo¨gliches Auftreten von Trogontherium cuvieri und Alces latifrons im letzten Interglazial der no¨rdlichen Oberrheinebene. Cranium 14, 2–10. Litt, T., 1994. Palaöo¨kologie, Palaöbotanik und Stratigraphie des Jungquarta¨rs im nordmitteleuropaïschen Tiefland. Dissertationes Botanicae 227, pp. 185. Litt. T., 2000. Waldland Mitteleuropa – die Megaherbivorentheorie aus palaöbotanischer Sicht. Berichte aus dem Bayerischen Landesamt fu¨r Wald und Forstwirtschaft 27, 49–64. Litt, T., Ellwanger D., Villinger, E., Wansa, S., 2005. Das Quarta¨r in der Stratigraphischen Tabelle von Deutschland 2002. Newsletter of Stratigraphy 41, 385–399. Mania D. (Ed.), 1997. Bilzingsleben V, Homo erectus – seine Kultur und Umwelt. Verlag AusbildungþWissen Bad Homburg Leipzig.

Mania, D., Thomae, M., 2006. Pleistoza¨nstratigraphie und Palaölithikum im mittleren Elbe-Saale-Gebiet. 73. Tagung der Arbeitsgemeinschaft Norddeutscher Geologen 6.–9. Juli 2006 Halle. Mania, D., Thomae, M., Litt, T., Weber, T., 1990. Neumark-Gro¨bern. Beitra¨ge zur Jagd des mittelpalaölithischen Menschen. Vero¨fftlichungen des Landesmuseums fu¨r Vorgeschichte in Halle 43, pp. 319. Pfeiffer, T., 1999. Sexualdimorphismus, Ontogenie und innerartliche Variabilita¨t der pleistoza¨nen Cervidenpopulationen von Dama dama geiselana PFEIFFER 1998 und Cervus elaphus L. 1758 (Cervidae, Mammalia) aus Neumark-Nord (Sachsen-Anhalt, Deutschland). Berliner geowissenschaftliche Abhandlungen E 30, 207–313. Sarntheim, M., Stremme, H.E., Mangini, A. 1986. The Holsteinian interglacial: Time-stratigraphic position and correlation to stable-isotope stratigraphy of deep-sea sediments. Quaternary Research 26, 283–298. Schreve, D.E., 2001. Mammalian evidence from the Middle Pleistocene fluvial sequences for complex environment change at the oxygene isotope substage level. Quaternary International 79, 65–74. Schreve, D.E., Bridgland, D.R., 2002. Correlation of Emglish and German Middle Pleistocene fluvial sequences based on mammalian biostratigraphy. Netherlands Journal of Geoscienes 81, 357–373. Stuart, A. J., 1981. A comparison of the Middel Pleistocene mammal faunas of Voigtstedt (Thuringia, GDR) and West Runton (Norfolk, England). Quarta¨rpalaöntologie 4, 155–163. Weber, T., Litt, T., Scha¨fer, D., 1996. Neue Untersuchungen zum a¨lteren Palaölithikum in Mitteldeutschland. Terra and Prehistoria – Beitra¨ge zur Ur- und Fru¨hgeschichte Mitteleuropas 9, 13–19.

30. MIS 5 to MIS 8 – Numerically Dated Palaeontological Cave Sites of Central Europe Wilfried Rosendahl1, Doris Do¨ppes2 and Stephan Kempe2 1

Reiss-Engelhorn-Museen, Zeughaus C5, D-68159 Mannheim, Germany Institut fu¨r Angewandte Geowissenschaften, Schnittspahnstr. 9, D-64287 Darmstadt, Germany

2

ABSTRACT Caves are among the most important sites preserving Quaternary fossils. Owing to the CaCO3-rich environment and the protection against erosion, even remains of early Pleistocene faunas are preserved in caves, while contemporaneous surface deposits have been lost. However, faunal remains cannot be linked to any interglacial or glacial period since no species exists which is characteristic of any specific period. Reliable dating of such remains is therefore required. This is now possible applying 230Th/U dating of speleothems. ESR dating of speleothems or 230Th/U dating of bones is, however, of disputable value. Dating of the base and top speleothem accretional layers permit assigning Pleistocene faunal remains to the MIS chronology. In this paper, we present for the first time an overview of all numerically dated palaeontological cave sites in Central Europe between MIS 5 and MIS 8. From twelve sites, a total of 31 strata were dated, most of them deposited during MIS 5; the rest belongs to MIS 6 and MIS 7; and only one sample representing MIS 8 provided reliable numerical dates. Numerically dated palaeontological cave sites older than MIS 8 are not known. 30.1 INTRODUCTION Caves are terrestrial depositories that may store a large variety of organic and inorganic remains. The latter may contain

important climatic and ecological information on Quaternary climate cycles. Caves are connected to the surface by joints, pits, sinkholes, dolines and horizontal passages. These serve to admit solid, fluid and gaseous matters directly or by way of intermediate deposits for final deposition, reflecting the changes of the environment above. Palaeontological remains are brought into the cave through a variety of processes. The excellent preservation of bones, molluscs and sometimes even other organic remains is due to the constant cave climate in which seasonal temperature changes are largely missing, and which has a rather constant humidity, wet in temperate climates and dry in desert settings. Even more important is the carbonate chemistry of the seepage waters and of the sediments which prevent dissolution of bones. Thus, only the soft tissue is lost due to bacterial decomposition, while the calcite saturated groundwater does not dissolve bones, teeth and shells. Conditions are also suitable for preserving pollen and charcoal (Bastin, 1978; Quinif and Bastin, 1994; McGarry and Caseldine, 2004) and even insect remains (Davis, 1999; Rosendahl and Kempe, 2002). Palaeontological remains from caves – specifically skeletal parts of Pleistocene large mammals – were among the first to be described scientifically (e.g. Esper, 1774; Rosenmu¨ller, 1794; Cuvier, 1805; Buckland, 1823). Their investigation marked the beginning of scientific cave research in Europe (e.g. Shaw, 1992). The dating of many species of Pleistocene mammals was established using bone deposits from caves or karst pits,

456

Wilfried Rosendahl, Doris Do¨ppes and Stephan Kempe

among them cave bear, cave hyena, cave lion, mammoth and woolly rhinoceros. Even though many mechanisms may cause deposition of bones in caves, it is generally possible to separate them into three groups (e.g. Zapfe, 1954; Rosendahl, 1995). The first group comprises the ‘cave dwellers’, animals which actively enter caves for temporary protection or to hibernate. These include a large variety of bat species and also large mammals such as the cave bear (Kurteń, 1976; Rabeder et al., 2000; Rosendahl et al., 2000), whose skeletons accumulated in large quantities in the so-called ‘bear caves’. The second group of fossils is due mainly to the activity of the cave hyena (Crocuta crocuta spelaea), a distinctly larger subspecies of the still-existing African spotted hyena (Crocuta crocuta). Hyenas can feed on bones; but in order to do so, they need to retreat to a protected site. Therefore, large bones, specifically those of the large Pleistocene herbivores, were carried individually into caves and partly consumed (e.g. Zapfe, 1939). ‘Hyena den caves’ therefore contain a variety of large, partly consumed, individual bones of those mammals living in the vicinity of the cave, but they never contain complete skeletons. Because cave hyenas are climate indifferent, the bone deposits have been accumulated both in interglacial and in glacial times. Also human activity can be a reason for bone accumulations in caves. Sometimes hyena and human occupation of the same cave have been observed (e.g. Weinbergho¨hlen bei Mauern, Koenigswald et al., 1974). ‘Prey animal deposits’ can also contain bones of small mammals when they derive from owls. At their daytime roosts owls and other nocturnal birds of prey regurgitate undigested skeletal remains, which can form sizeable bone deposits over the millennia. Such roosts are often found in cave entrances or in niches under overhanging ledges (abris). These deposits are most important for Quaternary palaeontology because small mammals are much better suited for the reconstruction of palaeo-

climatic and palaeoecological conditions than the occurrence of large mammals (Koenigswald, 1973; Heinrich, 1982, 1987). Small mammals occupy very characteristic environmental niches, and they react faster to climatic alterations than large mammals. The third group of bone deposits is formed accidentally. Either animals fall through fissures, pits and sink holes into the caves or their bones are secondarily transported into the caves by water or sediment. Thus, they may occur in deposits without much matrix as bone beds, or they may be part of a larger volume of clastic sediments. In this group, the distribution of species best reflects the fauna in the vicinity of the cave. Thus, fossils from caves can deliver important palaeoecological clues. Since many species are highly sensitive to climate and specific environments, their presence is a proxy for past environmental conditions. Cave fossils or fossil communities, however, cannot be used as dating tools a priori, since the same communities or animals may reoccur several times as the climate shifts between colder and warmer conditions. This is specially a problem during the MIS 6 and MIS 4. Nevertheless, certain mammal species do show a pronounced evolution throughout the Pleistocene and changes in certain physical characteristics, such as the structure of their teeth that can be used for an age proxy. When comparing characteristics of a certain species between different localities, it is often possible to deduce a relative temporal succession. Specifically certain small mammals, which in general show a fast evolutionary adaptation, have been found to be very valuable in this respect (Koenigswald, 1992). The Arvicolides for example play an important role for the biostratigraphic dating of the Middle Pleistocene since they evolved rapidly during this time interval (e.g. Koenigswald, 1973, 1992; Koenigswald and Heinrich, 1999). But even with these tools, only a relative age determination within the Pleistocene can be obtained (Lower Pleistocene, older or younger Middle Pleistocene or Upper Pleistocene)

Numerically Dated Palaeontological Cave Sites

(Koenigswald and Heinrich, 1996, 1999). Paralleling certain fossils with specific glacials or interglacials or Marine Isotope Stages (MIS) is only possible by numeric dating. Within the discussed time frame, only the 230 Th/U methods (alpha-spectrometry or TIMS) (e.g. Edwards et al., 1986/87) and with very strict limitations, the ESR method (e.g., Rink, 1997) can deliver useful age dates. In the following overview, we therefore summarize only those Central European cave sites (Fig. 30.1) for which numeric age data covering the period MIS 5 to 8 are available. The information regarding the respective sites was taken from the cited literature. A detailed critical discussion of the dating methods and the validity of the respective dates are, however, beyond the scope of this paper. It is necessary to point out that the individual dates obtained from bones or teeth must be viewed critically due to the open-system problem inherent to all samples which

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potentially suffered exchange with a fluid phase during their depositional history.

30.2 SITE DESCRIPTIONS 1 – Grotte Scladina Location: Belgium, Province of Namur, commune of Andenne, village of Sclayn Coordinates: 2 1893899E, 49 1794399N Altitude: 137.7 m a.s.l. (metre above sealevel) Local cave register number: Not registered Geographical position: The Scladina Cave is located in Sclayn, a village at the right bank of the Meuse Valley, ca. 60 km SE of Brussels. Site description: The 30 m long and 6 m high, tunnel-like cave was discovered in 1971; it is excavated scientifically since 1978 (Otte, 1992; Otte et al., 1998).

Fig. 30.1 Map of the reported Central European cave sites (numbers refer to the text).

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Stratigraphy: The main profile is 5.5 m thick and consists of intercalated clastic sediments and flowstone. It can be divided into 19 units (Otte et al., 1983). From top to bottom, these units are CC1, 36, 37, 38, 39 and 40, 1A, 1B, 2A, 2B, 3, CC4, 4A with CC14, 4B, 5, 6, 7A and 7B. Dating: Several flowstone layers were dated by the 230Th/U method in several laboratories (Gewelt et al., 1992). Four dates for unit 3 yielded an average age of 83 23 kyr. Layer CC4 was dated seven times. Two cores yielded mean ages of 114 23 and 110 14 kyr, respectively. This layer was also dated by TL yielding at its top 117:2 11 kyr and 122 11 kyr at its base (Debenham, 1998). From unit 5, a burnt artifact was also dated by TL and yielded 130 20 kyr (Huxtable and Aitken, 1992). According to magnetostratigraphical results, unit 3 coincides with MIS 5b, unit 4 with MIS 5c and unit 5 with MIS 5d (Ellwood et al., 2004). Fauna: Unit 3: Panthera leo spelaea, Crocuta crocuta spelaea, Canis lupus, Vulpes vulpes, Ursus spelaeus, Ursus arctos, Sus scrofa, Cervus elaphus, Rangifer tarandus, Dama dama, Capreolus capreolus, Bovinae, Capra ibex, Rupicapra rupicapra, Equus caballus, Hystrix cristata, Lepus sp. (Simonet, 1992). Microtus arvalis/agrestis, Microtus sp., Arvicola terrestris, Clethrionomys glareolus, Apodemus cf. sylvaticus, Talpa europaea, Citellus sp. (Cordy, 1992). Unit 4: Panthera leo spelaea, Panthera pardus, Felis sylvestris, Crocuta crocuta spelaea, Canis lupus, Cuon sp.?, Vulpes vulpes, Alopex lagopus, Ursus spelaeus, Ursus arctos, Ursus sp., Meles meles, Mustela putorius, Martes martes, Sus scrofa, Cervus elaphus, Rangifer tarandus, Dama dama, Capreolus capreolus, Bos primigenius, Bison priscus, Bovinae, Capra ibex, Rupicapra rupicapra, Mammuthus primigenius, Coelodonta antiquitatis, Equus caballus, Hystrix cristata, Lepus sp., Castor fiber (Simonet, 1992). Lagurus lagurus, Citellus sp., Dicrostonyx gulielmi, Microtus gregalis, Microtus oeconomus, Microtus arvalis/agrestis, Pitymys subterraneus, Microtus sp., Arvicola terrestris, Clethrionomys glareolus, Apodemus

cf. sylvaticus, Talpa europaea, Lemmus lemmus Chiropetra (Cordy, 1992). Unit 5: Panthera leo spelaea, Crocuta crocuta spelaea, Canis lupus, Ursus spelaeus, Ursus arctos, Ursus sp., Mustela foina, Sus scrofa, Cervus elaphus, Rangifer tarandus, Dama dama, Capreolus capreolus, Capra ibex, Rupicapra rupicapra, Mammuthus primigenius, Coelodonta antiquitatis, Equus caballus (Simonet, 1992). 2 – Einhornho¨hle Location: Germany, Lower Saxony, Harz, village of Scharzfeld Coordinates: 10 2491099E, 51 3891299N Altitude: 370 m a.s.l. Local cave register number: 4328/04 Geographical position: The Einhornho¨hle (Unicorn Cave) is located in the southern Harz Mountains near Scharzfeld, ca. 100 km SE of Hannover. Site description: The cave is more than 600 m long and consists of a string of large halls and domes connected by low passages. The first written record dates back to 1541, but the cave itself has been known much longer. It was an important source of bones sold as Unicornu fossile for medical purposes. Virchow (1872) began scientific excavations that continued until recently (e.g. Nielbock, 2003). The most recent excavation lasted from 1984 to 1988 (Nielbock, 2002). Stratigraphy: The 1.5-m thick standard profile from the ‘Weißer Saal’ encompasses the following layers: The ‘cave bear loam’ is sandwiched between a younger flowstone layer and a fossiliferous clay layer (Nielbock, 1987). In the Jacob-Friesen Passage, a more than 2-m thick profile yielded nine layers (0 – upper unit, H – lowest unit) (Nielbock, 1987). Dating: Cave bear bones of the ‘Weißer Saal’ yielded ESR dates between 95 and 104 kyr (Nielbock, 1987). Cave bear bones from the Jacob-Friesen Passage (units D to H) yielded 230Th/U dates of 126 þ10=9 kyr and 173 þ19=16 kyr (Wild et al., 1988). Fauna: ‘Weißer Saal‘; Canis lupus, Ursus spelaeus, Panthera leo spelaea (Nielbock, 1987).


Jacob-Friesen Passage (units D to H): Talpa europaea, Sorex araneus, Arvicola terrestris, Microtus nivalis, Microtus arvalis, Microtus agrestis, Microtus oeconomus, Canis lupus, Ursus spelaeus (Nielbock, 1987).

3 – Hunas Location: Germany, Bavaria, Fra¨nkische Alb, village of Hartmannshof Coordinates: 11 3294199E, 49 3091799N Altitude: 520 m a.s.l. Local cave register number: A 236 Geographical Position: In a limestone quarry near the Village of Hartmannshof, ca. 40 km E from Nuremberg. Site description: The cave was discovered in 1956 and investigated in the following eight years (Heller, 1983). The initial excavation opened the youngest part of an extensive stratigraphic sequence. In anticipation of the complete destruction of the site by the quarry, new excavations were started in 1983 and still go on. Stratigraphy: The cave is completely filled with layered sediments. The roof itself collapsed, covering the sediment-fill and sealing the cave entrance. About 12 m of sediments were investigated since 1983. The sediment stack can be divided into 22 layers. From top to bottom, these are the units A, B, C, D, E, F1, F2, G1, G2, G3, H, J, Koben, Kmitte, Kunten, Loben, Lmitte, Lunten, M, N, O and P (Rosendahl et al., 2006). Dating: In 2002, a flowstone layer was discovered at the base of the section (unit P). The layer is clearly connected with the cover sediment series without interruption. A stalagmite from this layer was mass spectrometrically dated (TIMS) by the 230Th/U method. The base yielded an age of 79 8 kyr and the top an age of 77 10 kyr (Rosendahl et al., 2005, 2006). Fauna: Unit O–P: Various genera and species of smaller and larger mammals occur, which have not been determined in detail as yet (Hilpert, personal communication).

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4 – Conturines-Ho¨hle Location: Italy, South Tyrol, Dolomites, village of S.Vigilio-Marebbe Coordinates: 11 599E, 46 349N Altitude: 2775 m a.s.l. Local cave register number: Not registered Geographical position: The ConturinesHo¨hle is located at the eastern slope of Piz dles Conturines, ca. 65 km E of Bozen. Site description: Fossil bone remains in the 160 m long Conturines-Ho¨hle are known since 1987 (Rabeder, 1991). Excavations were conducted from 1988 to 1990 (Rabeder, 1991), from 1996 to 1998 and in 2001. Stratigraphy: The floor of the upper parts of the cave is covered by thick flowstone. It is overlain in turn by fossiliferous dolomitic sand that is buried below large blocks (Rabeder, 1991). Dating: The basal flowstone is older than 350 kyr, beyond the range of dating by the 230 Th/U method (Frisia et al., 1993). The bone-bearing sands are much younger; the two oldest dates obtained by dating the bones are 87 5 kyr and 108 þ8=7 kyr (Withalm, 1995). Fauna: Marmota marmota, Ursus spelaeus, Panthera leo spelaea (Rabeder, 1991). 5 – Ramesch-Knochenho¨hle Location: Austria, Upper Austria, Totes Gebirge, village of Spital am Pyhrn Coordinates: 14 159E, 47 399N Altitude: 1960 m a.s.l. Local cave register number: 1636/8 Geographical position: The RameschKnochenho¨hle is located at the northface of the Ramesch, a peak in the Warscheneckgruppe/Totes Gebirge, W of Spital am Pyhrn, ca. 75 km S of Linz. Site description: Behind a wide entrance, a large, 30 m long hall with an almost level floor opens. The total length of the cave amounts to 310 m. First palaeontological excavations were conducted between 1979 and 1984 (Draxler et al., 1986).

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Stratigraphy: Undisturbed deposits are found only in the entrance hall (Draxler et al., 1986). Below a Holocene layer with gastropods (unit A) a typical cave loam with cave bear bones occurs (units B–E). Further down sterile loams fill the interstices between large blocks below dark sediments with cave bear remains, allochthonous pebbles (unit G) and fossil free pebblebearing sands (so-called ‘Augensteinsande’, unit H). Dating: Cave bear bones from unit G yielded 230Th/U dates between 117 þ11=10 kyr and 150 þ25=19 kyr (Draxler et al., 1986). Fauna: Unit G; Ursus spelaeus (Draxler et al., 1986).

6 – Schwabenreith-Ho¨hle ¨ tscherLocation: Austria, Lower Austria, O Lunzersee-Hochkar, village of Lunz am See Coordinates: 14 5893899E, 47 5093399N Altitude: 959 m a.s.l. Local cave register number: 1823/32 Geographical position: The Schwabenreith Ho¨hle is located S of the farm of Schwabenreith, W of the village of Lunz/ See, and ca. 120 km SE of Linz. Site description: The 134-m long cave has two nearly horizontal passages. The cave has been known locally for a long time but was just investigated only at the end of the 1960s (Hartmann and Hartmann, 1969). Excavations lasted from 1990 to 2000 (Fladerer, 1992; Pacher, 2000). Stratigraphy: The terminal hall yielded most of the fossils (Site 2) where stalagmites sit directly on the rock floor. Above them a fossil-free loamy sand (unit 6) was deposited followed by a 1.3-m thick fossiliferous layer (unit 5), which is covered by a layer of flowstone (unit 2). Dating: The basal and top (unit 2) flowstone layers yielded 230Th/U ages of 116 5 kyr and 78 þ 30=23 kyr, respectively (Frank and Rabeder, 1997b). Fauna: Ursus spelaeus (Pacher, 2000).

7 – Herdengelho¨hle Location: Austria, Lower Austria, ¨ tscher-Lunzersee-Hochkar, village of O Lunz am See Coordinates: 14 5893899E, 47 5092599N Altitude: 878 m a.s.l. Local cave register number: 1823/4 Geographical position: The cave is situated SW of the farm Herdengel, W of the village of Lunz am See, ca. 120 km SE of Linz. Site description: The cave is 129 m long. Digs conducted in 1927 and 1928 remained without success but in 1935 fossils were discovered (Abrahamczik, 1936). Modern scientific excavations were conducted from 1983 to 1989 (Nagel and Rabeder, 1991; Frank and Rabeder, 1997a). Stratigraphy: The 8-m thick sediments contain six layers. The lowest layer (down to 750 cm) is a sterile, fine sand covered by flowstone. Unit 1 (380–430 cm) above it is a black yellow loam which contained bones stained black. Unit 2 (360–380 cm), a flowstone layer, shows well-developed stalagmites. Unit 3 (330–360 cm) and unit 4 (300–330 cm) contained partially-to-well preserved brown-stained bones. Unit 5 (280–300 cm) is composed of blocks and reddish loam. Unit 6 (200–280 cm) forms a 2-m thick cover composed of fossil-free, light yellow loam (Leitner-Wild et al., 1994). Dating: The flowstone layer of unit 2 was 230 Th/U dated to 111þ11=10 kyr. The cave bear bones of unit 1 date back to the period from 135 þ11=10 kyr to 127 7 kyr (Leitner-Wild et al., 1994). Fauna: Unit 1; Ursus spelaeus (Frank and Rabeder, 1997a). 8 – Repolustho¨hle Location: Austria, Styria, Mittelsteirischer Karst, village of Frohnleiten Coordinates: 15 2095199E, 47 1893599N Altitude: 525 m a.s.l. Local cave register number: 2837/1 Geographical position: The Repolustho¨hle is located in the valley of Badl, ca. 20 km N of Graz.


Site description: The cave was discovered in 1910 and is 66 m long. It ends in a natural filled pit. Excavations were conducted between 1947 and 1955 (Mottl, 1951; Mottl and Murban, 1955). The sediments of the pit were investigated by Temmel (1996), and the material of old excavations was revised by Fu¨rnholzer (1997). Stratigraphy: The sedimentary profile in the pit (from the bottom to the top) shows loam and clay followed by rust-coloured phosphate-rich sediments with manganese streaks, grey sands and a rust-coloured phosphate-rich soil (Mottl and Murban, 1955). Dating: A cave bear bone from the lowest layers in the pit yielded a 230Th/U age of 230 þ 13=12 kyr (Fu¨rnholzer, 1997). Fauna: Lower rust-coloured phosphaterich sediment; Aves, Talpa europaea, Sorex cf. araneus, Myotis bechsteini, Plecotus auritus, Marmota marmota, Spermophilus cf. citellus, Cricetus major, Apodemus sylvaticus, Apodemus flavicollis, Clethrionomys glareolus, Arvicola hunasensis, Microtus arvalis, Hystrix cf. vinogradovi, Lepus sp., Canis lupus ?, Canis mosbachensis, Vulpes vulpes, Cuon alpinus ssp., Ursus arctos, Ursus deningeri, Martes martes, Meles meles, Mustela nivalis, Putorius sp., Felis silvestris, Lynx lynx, Panthera pardus, Panthera leo spelaea, Sus scrofa, Cervus elaphus, Megaloceros giganteus, Capreolus capreolus, Rangifer tarandus, Bison priscus, Rupicapra rupicapra, Capra ibex, Elephantidae indet (Rabeder and Temmel, 1997). 9 – Divje babe I Location: Slovenia, Innerkrain (Notranjska), village of Reka Coordinates: 13 549E, 46 019N Altitude: 450 m a.s.l. Local cave register number: 812 Geographical position: The Divje Babe I site is located in the Idrija valley near the village of Reka, at the western slope of the Sˇebrelje plateau, ca. 60 km W of Ljubljana. Site description: The cave is approximately 45 m long and 45 m wide. Systematic excavations lasted from 1980 to 1986 (Turk

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et al., 1989) and from 1989 to 1995 (Turk et al., 2002). Stratigraphy: This Upper Pleistocene site contains a long stratigraphic sequence spanning the period from approximately 120 to 35 kyr (Turk et al., 2001). The 11.5-m thick profile is divided into 21 units; the oldest is unit 21 (Turk et al., 2002). Dating: 230Th/U dating of a bear bone yielded 81 10 kyr (unit 20) and 84 7 kyr for unit 19 (Nelson, 1997). Fauna: Unit 20–17; Ursus spelaeus (99%) (Turk et al., 1989). 10 – Vindija Location: Croatia, Hrvatsko Zagorje, village of Ivanec Coordinates: 16 2093899E, 46 199N Altitude: 275 m a.s.l. Local cave register number: Not registered Geographical position: Vindija is located on the southwest side of the Kriznjak Peak, 55 km N of Zagreb. Site description: The cave chamber is approximately 50 m long, 28 m wide and more than 10 m in height. Vindija was first mentioned in 1878, and excavation was started in 1928 (Malez, 1979). Extensive archaeological and palaeontological excavations lasted from 1974 to 1986 and from 1993 to 1994 (Malez et al., 1984; Karavanic´, 1995). Stratigraphy: The 12-m thick stack of sediments in Vindija cave can be divided into 13 layers designated from unit A (youngest) to unit M (oldest). Dating: 230Th/U dates from cave bear bones of unit J range from 156 2 kyr to 196 þ20=15 kyr and of the underlying unit K from 150þ16=13 kyr to 212þ17=13 kyr (Wild et al., 2001). Fauna: Unit J: Marmota marmota, Canis lupus, Cuon alpinus europaeus, Ursus spelaeus, Panthera leo spelaea, Cervus elaphus, Capra ibex and Bovidae indet. (Malez and Ullrich, 1982). Unit K: Canis lupus, Ursus spelaeus, Panthera leo spelaea, Panthera pardus, Crocuta crocuta spelaea, Dicerorhinus kirchenbergensis,

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Sus scrofa, Megaloceros giganteus, Cervus elaphus, Dama dama, Capreolus capreolus, Bos primigenius (Malez and Ullrich, 1982). 11 – Krapina Location: Croatia, Hrvatsko Zagorje, village of Krapina Coordinates: 15 529E, 46 109N Altitude: 120 m a.s.l. Local cave register number: Not registered Geographical position: Krapina (Husˇnjakovo rock shelter) is located on the western side of the Husˇnjak hill, W of Krapina, 42 km NW of Zagreb. Site description: The rock shelter was ca. 12 m high. The site was excavated from 1899 to 1905 (Gorjanovic´-Kramberger, 1906). Stratigraphy: The 12-m thick stack of sediments in Krapina can be divided into 10 layers designated unit I (oldest) to unit 9 (youngest). Dating: The age of tooth enamel of hominids from unit 9 yielded a 230Th/U date of 113 10 kyr and an ESR date of 87 7 kyr. The ages of teeth from units 1–8 were indistinguishable, with a mean of 130 10 kyr (Rink et al., 1995). Tooth enamel of hominids from unit 9 to 6 and unit 1 was dated by U-series dates and ESR (Rink et al., 1995). Fauna: Unit 9: Lepus sp., Marmota marmota, Canis lupus, Ursus spelaeus, Lynx lynx, Cervus elaphus, Rupicapra rupicapra (Malez, 1970) Units 7 and 8: Aves, Amphibia, Castor fiber, Myoxus glis, Canis lupus, Ursus spelaeus, Ursus arctos priscus, Mustela putorius, Martes martes, Felis silvestris, Panthera pardus, Sus scrofa, Cervus elaphus, Capreolus capreolus, Bos primigenius, Stephanorhinus kirchbergensis, Homo neanderthalensis (Malez, 1978). Units 5 and 6: Cricetus cricetus, Canis lupus, Vulpes vulpes, Ursus spelaeus, Ursus arctos priscus, Mustela cf. eversmanni, Lynx lynx, Cervus elaphus, Alces alces, Bison priscus, Equus cf. germanicus, Homo neanderthalensis (Malez, 1978). Unit 1: Castor fiber, Panthera pardus, Stephanorhinus kirchbergensis, Homo neanderthalensis (Malez, 1970).

12 – Bisńik Jaskinia Location: Poland, Province of Ło´dz´, Krakow Cze˛stochowa Uplands, village of Pilica Coordinates: 19 559E, 50 289N Altitude: 395 m a.s.l. Local cave register number: Not registered Geographical position: The Bisńik Jaskinia is located near the village of Pilica in the central part of the NiegowonickoSmolen´skie Hills, ca. 50 km N of Cracow. Site description: The cave is about 73 m long. It consists of a rock shelter and a cave proper connected to it. Excavation lasted from 1991 to 2000 (Mirosław-Grabowska, 2002). Stratigraphy: The more than 7-m thick clastic sediments can be divided into 18 layers. The lowest series consists of layers 8 to 18 (Mirosław-Grabowska, 2002). Dating: 230Th/U dating of bones from layers 12 and 13 yielded an age range of 115–128 kyr, for layer 14 from 128 to 200 kyr, for layer 15 from 200 to 250 kyr and for layers 16 and 17 from 250 to 270 kyr (MirosławGrabowska, 2002). Fauna: Layers 12, 13: Ursus spelaeus, Cervus elaphus, Capreolus capreolus, Rangifer tarandus, Bos primigenius, Equus caballus, Clethrionomys glareolus (Mirosław-Grabowska, 2002). Layer 14: Canis lupus, Ursus spelaeus, other carnivores, Sus scrofa, Megaloceros giganteus, Cervus elaphus, Rangifer tarandus, Bison priscus, Equus caballus, Arvicola terrestris, Microtus oeconomus, Clethrionomys glareolus (Mirosław-Grabowska, 2002). Layer 15: Canis lupus, Vulpes vulpes, Ursus spelaeus, other carnivores, Sus scrofa, Cervus elaphus, Capreolus capreolus, Rangifer tarandus, Bison priscus, Equus caballus, Arvicola terrestris, Microtus oeconomus, Apodemus sylvaticus, Clethrionomys glareolus (Mirosław-Grabowska, 2002). Layers 16, 17: Ursus spelaeus, Crocuta crocuta spelaea, Vulpes vulpes, other carnivores, Cervus elaphus, Megaloceros giganteus, Capreolus capreolus, Alces alces, Rangifer tarandus, Bison priscus, Coelodonta antiquitatis, Equus caballus (Mirosław-Grabowska, 2002).


30.3 CONCLUSIONS In spite of the fact that numerous palaeontologically important cave sites are known in Central Europe which may be dated into the time period MIS 5 to MIS 9 (e.g. Koenigswald and Heinrich, 1996, 1999; Do¨ppes and Rabeder, 1997), we found only twelve sites in the literature for which numerical dates have

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been published. A total of 31 layers have been dated in these sites (Fig. 30.2). Most of them yielded dates corresponding to MIS 5, a few represent MIS 6 and 7 and only one can be correlated with MIS 8. Older numerically dated sites have not yet been found. More than half of the faunal strata have been dated because they are part of archaeologically important sites. Numerical dates of purely

Fig. 30.2 Timetable and assignment of cave deposits to the MIS chronology. Dot date without standard deviation, dot and line date with a standard deviation, solid line time range of strata with several dates. Explanation of abbreviations: BC Bisńik Jaskinia (12 – number of site description), Cu Conturines-Ho¨hle (4), DB ¼ Divje babe I (9), Hd Herdengelho¨hle (7), Hu Hunas (3), Kr Krapina (11), Re Repolustho¨hle (8), Rk Ramesch-Knochenho¨hle (5), SC Grotte Scladina (1), Sw Schwabenreith-Ho¨hle (6), Uc Einhornho¨hle (2), Vi Vindija (10).

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palaeontological sites are rare even though the archive of caves and their rich bone beds offer enough potential for a systematic study. There are several reasons why this has not been done yet: First of all, cave sites are not easily accessible (compared to open air sites); second, the mechanisms of deposition are complicated and can only be understood by investigating the genesis of the respective cave and its sediments in general. Cave studies therefore require additional knowledge (i.e. speleology) before the specific site can be interpreted correctly. The most important point, why so few dates exist is related to the dating itself. Beyond the reach of the 14C method, bones, as mentioned above, can only be dated by the 230Th/U or ESR methods. The 230Th/U method is the one most commonly used. But ESR dates of bones are highly problematic and should not be used (e.g. the dates of the Einhornho¨hle, no. 2). Only ESR dates of tooth enamel appear to be correct (e.g. Krapina, no. 11). TL dating, used for speleothems (e.g. Grotte Scladina, no.1) exclusively, is methodologically also of doubtful quality, and TL dates should today be regarded with caution and their usage should be discontinued. Even the 230Th/U dating of bones is methodologically problematic due to the fact that bones very often prove to be open systems (for a discussion see Bischoff et al., 1995). The unusual standard deviations of the dates of unit J of Vindija (no. 10; Wild et al., 2001) may be caused by exactly this open-system problem. Therefore, it is essential that prior to the dating both the excavator and the dating geochemist discuss the chronostratigraphy, palaeoecology and palaeoclimatology of a site in detail. Even though, there are isotopists who view all 230Th/U bone dates critically and suggest they be discarded all together (e.g. Geyh, 2005). The 230Th/U-dating technique of speleothems has been improved substantially, resulting in more reliable results since the 1990s. These have been used for the reconstruction of Middle to Upper Pleistocene climate and environment (e.g. Winograd

et al., 1992; Kempe et al., 2002; Genty et al., 2003; Holzka¨mper et al., 2005), the dating of bones is lagging behind. Only a few laboratories (e.g. Vienna and Warsaw) are currently applying it. It would therefore be profitable if the technique of dating bones with 230Th/U could also be improved. Interesting suggestions in this direction have been made by Hercman and Gorka (2000), Pike et al. (2002) and Eggins et al. (2005). Hoffmann and Mangini (2003) also describe an interesting method to date teeth and perhaps bones from open systems. Even though speleothem dating with the TIMS 230Th/U method is not entirely free of methodological problems, TIMS speleothem dates are the best dates available today to establish cave-based chronologies. Flowstone layers above or below the faunal strata can thus be dated, bracketing the ages of the bones (e.g. Schwabenreith-Ho¨hle, no. 6). Those sites, which do not have speleothem supported age models, should therefore be revisited, and additional samples should be dated to give the currently available bone dates further credibility. For climatic and ecological investigations of speleological faunas, we should therefore target those cave sites which can be dated via speleothems. Additionally, bones could be dated with TIMS 230Th/U and teeth with ESR in order to advance dating techniques in general. Since the open-system problem induces a substantial inaccuracy regarding dates of bones and teeth, a critical assessment of specific faunal assemblages with respect to their exact stratigraphical position remains difficult. In addition, numerical dates have a certain standard deviation caused by methodological problems. This deviation can be substantial with the consequence that only two of the faunal assemblages suitable for ecological discussion can be attributed to either a glacial or an interglacial. These two faunas are layers 12–13 of the Bisnik Jaskinia (attributed to MIS 5e) and the fauna recovered from the pit of the Repolust Cave (attributed to MIS 7).


MIS 5 Insectivora Sorex araneus Talpa europaea Chiroptera Myotis bechsteinii Plecotus auritus Lagomorpha Lepus sp. Rodentia Marmota marmota Spermophilus sp. Spermophilus citellus Castor fiber Cricetus cricetus Cricetus major Lemmus lemmus Dicrostonyx gulielmi Clethrionomys glareolus Lagurus lagurus Arvicola terrestris Arvicola hunasensis Microtus sp. Microtus agrestis Microtus arvalis Microtus agrestis /arvalis Microtus gregalis Microtus oeconomus Microtus subterraneus Chionomys nivalis Apodemus flavicollis Apodemus sylvaticus Glis glis Hystrix cristata Hystrix cf. vinogradovi Carnivora Canis aureus Canis lupus Canis mosbachensis Cuon sp.? Cuon alpinus Alopex lagopus Vulpes vulpes Ursus sp. Ursus arctos Ursus spelaeus Ursus deningeri Mustela sp. Mustela eversmanii Mustela nivalis Mustela putorius Martes foina

Common shrew Common mole

+ +

465

MIS 6

MIS 7

+ +

cf. +

Bechstein’s bat Brown long-eared bat

+ +

Hare

+

Alpine marmot Suslik European suslik Eurasian beaver Common hamster Giant hamster Norway lemming Collared lemming Bank vole Steppe lemming Nothern water vole Water vole Vole Field vole Common vole Field/common vole Narrow-headed vole Tundra vole Pine vole Snow vole Yellow-necked mouse Wood mouse Edible dormouse Crested porcupine Porcupine

+ +

+

+ +

+ +

Golden jackal Grey wolf Tautavel wolf Dhole Indian dhole Arctic fox Red fox Bear Brown bear Cave bear Deninger bear Weasels Steppe polecat Weasel Western polecat Beech marten; Stone marten

MIS 8

+ + cf.

+ + + + + + + + + + + + +

+

+

+

+ +

+

+

+

+ cf. + +

+ + + cf.

+

+

? +

+

+

+ + + + + +

+ + +

cf.

cf.

+ +

+ +

+ + + + +

+

+

+

Fig. 30.3 (table) Faunal distribution of the reported cave sites from MIS 5 to MIS 8 without critical reflection on the numerical dates.


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Martes martes Meles meles Felis silvestris Panthera leo spelaea Panthera pardus Lynx lynx Crocuta crocuta spelaea Artiodactyla Sus scrofa Dama dama Cervus elaphus Megaloceros giganteus Alces alces Rangifer tarandus Capreolus capreolus Bovidae indet. Bovinae Bos primigenius Bison priscus Rupicapra rupicapra Capra ibex Proboscidae Elephantidae indet. Mammuthus primigenius Perissodactyla Equus caballus Equus cf. germanicus Coelodonta antiquitatis

Pine marten Badger European wild cat Cave lion/Fig. 30.4, a Leopard European lynx Cave hyena Wild boar Fallow deer Red deer Giant deer Elk, Moose Reindeer Roe deer Bovid Bovid Aurochs Steppe wisent Alpine chamois Alpine ibex

MIS 5 + + + + + + +

MIS 6 +

+ + +

+ + + + + + + +

+

+ + + + +

+ + + +

+ + + +

+

+

+ + + +

+ + + +

+

+

+ + +

Wood elephant?/Fig. 30.4, b Woolly mammoth

Wild horse Horse Woolly rhinoceros/Fig. 30.4, c Stephanorhinus kirchbergensis Merck’s rhinoceros Primate Homo neanderthalensis Neanderthalian/Fig. 30.4, d

+ + + + +

MIS 7 + + + + + + +

+ + + +

MIS 8

+

+ + + + +

+

+

+

+ +

+

Fig. 30.3 Continued

In case of Bisnik Jaskinia, only the Rangifer tarandus component is in contrast to its Eemian age, but the faunal remains of the Repolust Cave combine both glacial and interglacial species, i.e. Rangifer tarandus and Megaloceros giganteus occur together with Capreolus capreolus and Sus scrofa. Thus, the presence of cold climate species is in contrast to the numerical interglacial date (Fig. 30.3, 30.4). The accuracy of the numerical dates of all other sites and layers do not permit to attribute the faunas into a specific glacial or interglacial. Thus the faunas cannot be evaluated regarding their ecological and climatic character. Contradicting occurrences of glacial and interglacial faunal elements cannot be resolved as long as the numerical dates allow for both possibilities.

In conclusion, the now available numerical dates of palaeontological sites in Central Europe do not allow – with the exception of Bisnik Jaskinia and the Repolustho¨hle Cave – a critical discussion of their faunal assemblages as to their ecological climatic distribution. But even those two sites are not without contradicting faunal elements, and it remains doubtful if they represent either glacial or interglacial faunas. In spite of all these problems, palaeontological cave sites represent a rich archive that can deliver important contributions to the reconstruction of the Middle and Upper Pleistocene palaeoclimate of Central Europe, provided many additional dates can be obtained to verify results obtained from other terrestrial archives.


(a)

(b)

(c)

(d)

467

Fig. 30.4 Lifestile reconstructions of different faunal elements named in Fig. 30.3: a=Panthera leo spelaea (climate indifferent); b = Elephas antiquus (interglacial); c = Coelodonta antiquus (glacial) d = Homo neanderthalensis (climate indifferent). All photographs W. Rosendahl, Mannheim.

ACKNOWLEDGEMENTS The authors thank the reviewers for critical and helpful remarks and Dr. M.S. Werner (Hilo/Hawaii) for editorial suggestions. REFERENCES Abrahamczik, W., 1936. Karsterscheinungen in der Umgebung von Lunz am See (mit besonderer Beru¨cksichtigung der Ho¨hlen). Unpublished Ph. D., University Vienna, 100 pp.

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Numerically Dated Palaeontological Cave Sites Huxtable, J., Aitken, M.J., 1992. Thermoluminescence dating of burned flint and stalagmitic calcite. E´tudes et recherches archeólogiques de l’Universite´ de Lie`ge 27, 175–178. Karavanic´, I., 1995. Upper Paleolithic occupation levels and late-occurring Neandertal at Vindija Cave (Croatia) in the context of Central Europe and the Balkans. Journal of Anthropological Research 51, 9–35. Kempe, S., Rosendahl, W., Wiegand, B., Eisenhauer, A., 2002. New Speleothem Datation from Caves in Germany and their importance for the Middleand Upper-Pleistocene Climate Reconstruction. Acta Geologica Polonica 52, 55–61. Koenigswald, W. v., 1973. Vera¨nderungen in der Kleinsaügerfauna von Mitteleuropa zwischen Cromer und Eem (Pleistoza¨n). Eiszeitalter und Gegenwart 23/24, 159–167. ¨ kologie und BiostratiKoenigswald, W. v., 1992. Zur O graphie der beiden pleistoza¨nen Faunen von Mauer bei Heidelberg. In: Beinhauer, K.W., Wagner, G.A. (Eds.), Schichten von Mauer – 85 Jahre Homo erectus heidelbergensis, Braus Mannheim, 101–110. Koenigswald, W. v., Heinrich, W.-D., 1996. Kurze Charakterisierung der Vera¨nderungen in der Saügerfauna des Jungquarta¨rs in Mitteleuropa. Tu¨binger Monographie zur Urgeschichte 11, 441 pp. Koenigswald, W. v., Heinrich, W.-D., 1999. Mittelpleistoza¨ne Saügetierfaunen aus Mitteleuropa – der Versuch einer biostratigraphischen Zuordnung. Kaupia 9, 53–112. Koenigswald, W. v., Mu¨ller-Beck, H., Pressmar, E., 1974. Die Archaölogie und Palaöntologie in den Weinbergho¨hlen bei Mauern (Bayern), Grabungen 1937–1967. Archaeologica Venatoria 3, 1–152. Kurteń, B., 1976. The Cave Bear Story, Life and Death of a Vanished Animal. Columbia University Press, 163 pp. Leitner-Wild, E., Rabeder, G., Steffan, I., 1994. Determination of the evolutionary mode of Austrian alpine cave bears by uranium series dating. Historical Biology 7, 97–104. Malez, M., 1970. Die Ergebnisse der Revision der pleistoza¨nen Fauna aus Krapina. Krapina 1899– 1969, 45–56. Malez, M., 1978. Stratigraphische, palaöfaunistische und palaölithische Verha¨ltnisse des Fundortes Krapina. Krapinski pracovjek I evolucija hominida, Rad Jugoslavenske akademije znanosti i umjetnosti 61–102. Malez, M., 1979. Paleolitsko i mezolitsko doba u Hrvatskoj, Praistorija jugoslavenskih zemalja I, pp. 227–295. Malez, M., Simunis, An., Simunis, Al., 1984. Geoloki, sedimentoloki i paleoklimatski odnesi spilje Vindje i blize okolice. Rad Jugoslavenske akademije znanosti i umjetnosti 411, 231–264.

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31. The Last and the Penultimate Interglacial as Recorded by Speleothems From a Climatically Sensitive High-Elevation Cave Site in the Alps Christoph Spo¨tl1, Steffen Holzka¨mper2 and Augusto Mangini3 1

Institut fu¨r Geologie und Palaöntologie, Leopold-Franzens-Universita¨t Innsbruck, Innrain 52, 6020 Innsbruck, Austria 2 Department of Physical Geography and Quaternary Geology, Stockholm University, 10691 Stockholm, Sweden 3 Forschungsstelle Radiometrie, Heidelberger Akademie der Wissenschaften, Im Neuenheimer Feld 229, 69120 Heidelberg, Germany

ABSTRACT Within the greater Alpine region, absolutely dated climate records of the penultimate and the last interglacial are exceptionally rare. Speleothems offer an important and still underutilized source of information about the timing and duration of warm periods during the Middle and Upper Quaternary. The focal point of intense research is Spannagel Cave, a large high-altitude (ca. 2200 to 2500 m a.s.l.) cave system in the Zillertal Alps of Austria. The presently low (1.4 to 2.5 C) cave temperature provides a natural threshold for speleothem growth, i.e. the cave acts as a climatically sensitive archive. U-series dating of calcite speleothems, facilitated by exceptionally high U content in combination with highresolution stable isotope analyses allow identifying warm climate periods. Calcite growth at 236 to 229 kyr, 211 to 206 kyr and 199 to 192 kyr is in good accordance with U-series dated sea-level records and marine sediments whose chronology was tuned to orbital parameters. Oxygen isotope data show that the climate in the Alps was consistently cooler during the penultimate interglacial than during the last interglacial. Carbon isotope data also show a major difference between the two interglacials: while alpine soil and vegetation was apparently well developed during the last interglacial (and similar as today), high C

isotope values testify the lack of pedogenic C input during the penultimate warm periods. Accordingly, the area above the cave was either barren or – more likely – covered by a warm-based glacier. Previously regarded as evidence of ice-free conditions early during the penultimate deglaciation speleothem deposition at 136 kyr is now seen as an indication of a major change of the glacier’s thermal state most likely as a result of the collapse of ice-stream network at the end of the penultimate glacial maximum. Following a return to stadial conditions not conducive to speleothem formation and marked by a hiatus in speleothem growth, fully interglacial conditions did not commence until 130 kyr and prevailed until 119 kyr.

31.1 INTRODUCTION The Alps are probably the most thoroughly studied mountain range on Earth as far as climate history is concerned. This is particularly true for the instrumental period (i.e. since the middle of the eighteenth century). Recent studies, for instance, have shown that the Alps warmed by about 1.5 C since the end of the ‘Little Ice Age’, which is more than twice as much as the mean Northern Hemisphere warming and similar to the temperature trend observed in the Arctic

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(Bo¨hm et al., 2001, personal communication 2005; Moritz et al., 2002; Arctic Climate Impact Assessment, 2004; Climate Research Unit, 2004). The reason for this strong regional increase in temperature during the past 150 years remains to be fully understood. In the light of record high temperatures in the past two decades (e.g. 1998, 2002, 2003) – absent in even the longest instrumental records – a long-term perspective of the natural climate variability in the Alps is sought. Recent studies on wood fragments released by receding alpine glaciers demonstrate that during the first part of the Holocene, the glaciated area of the Alps was repeatedly substantially smaller than today. These studies also suggest a pattern of repeated rises and falls of the equilibrium line altitude during the Holocene (Hormes et al., 2001; Nicolussi and Patzelt, 2000a, 2000b) that may possibly be related to the millennial scale climate variability of the North Atlantic (Bond et al., 2001; Broecker, 2001). Previous interglacials can provide additional insights into climate and environmental changes in the Alps. Until now, however, climatic records from pre-Holocene interglacials are extremely sparse and fragmentary within the Alps, owing to the destructive action of ice during the glacials. Therefore, our current understanding of climate change during the last interglacial relies on often nondated palaeovegetation data from low-lying sites in the foreland of the Alps. Hardly any reliable record is available from an interglacial prior to the last interglacial, and even the latter successions are mostly dated by means of biostratigraphic correlation only (e.g. Gru¨ger, 1979; Drescher-Schneider, 2000; Mu¨ller et al., 2003; Mu¨ller and Kukla, 2004). In recent years, a new type of palaeoclimate information has been retrieved from the shallow subsurface from speleothems in caves. These secondary carbonates have seldom been affected by erosion processes and may contain a wealth of palaeoclimate proxy information that can be precisely dated using state-of-the-art mass spectrometric U-series methods (see

Dorale, 2004; White, 2004 for recent reviews of this subject). Speleothem deposition requires the presence of groundwater supersaturated with respect to calcite. In regions such as the Alps, the limiting factor for speleothem formation is temperature. Consequently, speleothem growth intervals obtained from studies of presently cool alpine caves provide important temporal constraints on the timing of warm climate periods in the past. This article provides an overview of recent efforts of our group working in high-alpine cave sites to retrieve, analyse and evaluate this palaeoclimate information. Focusing on interglacial growth periods from Spannagel Cave in Tyrol, it considers and re-examines both previously published work and new, hitherto unpublished palaeoclimatic proxy data. 31.2 CAVE SETTING Spannagel Cave is located in the Central Alps of Austria and comprises a network of slightly more than 10 km of passages and short shafts located between 2524 and 2195 m (Fig. 31.1). It is the largest out of a series of more than 30 caves that developed within the Jurassic calcite marble that forms a tectonically deformed, 20-m thick slab dipping towards N and NNW beneath granitic gneiss. A crucial aspect is the proximity of the cave to the Hintertux Glacier (Fig. 31.1). Developed partly beneath a broad ridge separating two adjacent glacially shaped valleys about half of the cave system was in a subglacial position as recent as during the ‘Little Ice Age’, which was the most extensive glacier advance during the Holocene in the Alps (Maisch et al., 2000; Hormes et al., 2001). Sharp crested, poorly vegetated lateral moraines mark the 1850 advance of the Hintertux Glacier (Fig. 31.1). During periods when the ice extent was larger than this advance, most if not all of the cave was buried beneath

The Last and the Penultimate Interglacial

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(a)

Karstifiable carbonate rocks in Austria

VIENNA

LINZ

SALZBURG

BREGENZ

INNSBRUCK GRAZ

KLAGENFURT

Till

(b)

500 m

Marble Granitic gneiss Paragneiss Glacier tongue during ‘Little Ice Age’

N

Spannagel Cave

Spannagel hut

Hintertux Glacier

(c) Lateral moraine ridge (‘Little Ice Age’ advance) 100 m

Spannagel hut Cave entrance

Fig. 31.1 Simplified geologic map of the area near Spannagel Cave, western Zillertal Alps, Tyrol. The plan view of the cave is superimposed on the map. Other, smaller caves present in this region are omitted for clarity. The location of the cross section (C) below is indicated by the two green arrows in B. The cave in bound to the thin slab of marble. Note different scales.


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ice, reaching a maximum thickness of ca. 250 m during the last glacial maximum (van Husen, 1987). The present-day air temperature in the cave is constant during the year and slightly above the freezing point (Fig. 31.2) allowing water–rock interactions and (slow) formation of speleothems. Contrary to many other high-elevation caves in the Alps – which contain none or ancient speleothems only– several passages of Spannagel Cave provide evidence of modern dripstone formation, such as soda straws, stalactites, stalagmites and flowstones (Figs. 31.2 and 31.3). According to multiannual cave water monitoring dripwater feeding, speleothems is thermodynamically supersaturated with respect to calcite (either year around or seasonally; Spo¨tl, unpublished data), and 230Th/U dates demonstrate a Holocene age of these commonly active speleothems (Fig. 31.2). Other passages of Spannagel Cave contain speleothems – locally abundant – that are clearly inactive and partially broken and/or

corroded. 230Th/U dates document several episodes of speleothem growth during the past few hundred thousand years, including some dates exceeding the limit of the 230Th/ U method (Spo¨tl et al., 2004; Fig. 31.2). We identified a set of flowstones and stalagmites that grew during the penultimate interglacial (marine isotope stage 7, MIS 7) and the last interglacial (MIS 5.5). These samples were collected from different parts of the large cave system (Fig. 31.2). The dates allow validating previous findings based on isotope analyses of single speleothem samples. 31.3 METHODS The chronology of speleothem samples from Spannagel Cave was established by 230Th/U TIMS dating at Heidelberg University. For analytical details, see Frank et al. (2000). As a result of spike recalibration during the second half of 2004 and calibration against the ‘HU-1’ uraninite standard solution SPA11,52 SPA50,51

N 1.5

1.8

100 m

SPA129 1.9

Speleothems

1.6

Holocene and modern Last interglacial Penultimate interglacial

2.5

Entrance 1.4

1.7

SPA59

Fig. 31.2 Plan view of the cave network. Symbols indicate 230Th/U dated flowstone and stalagmite samples of Holocene, last interglacial and penultimate interglacial age. Samples examined in this study are labelled. Not shown are speleothems older than MIS 7 as well as samples formed during MIS 3. Yellow labels refer to cave air temperatures (constant year around) in degrees Celsius based on multiannual monitoring.


475

31.4 THE SPELEOTHEM RECORD

Fig. 31.3 This currently dry passage of Spannagel Cave shows evidence of Holocene speleothem formation (white carbonate deposits on ceiling and walls) and of former presence of a cave stream which left a sandy rubble blanket on the passage floor.

(assuming that uraninite is at secular equilibrium for the 230Th-234U-238U sequence), previously published 230Th/U dates were systematically too young. Revised 230Th/U dates of these results published by Spo¨tl et al. (2002) and Holzka¨mper et al. (2004) are provided by Holzka¨mper et al. (2005). Additional hitherto unpublished 230Th/U dates are presented in this article. A micromill technique was used to obtain continuous, high-resolution stable carbon and oxygen isotope transects along 230Th/Udated speleothem samples. The spatial resolution was 0.10–0.15 mm, about one order of magnitude higher than in our previous studies. The isotopic compositions were measured using a DeltaplusXL mass spectrometer equipped with an automated carbonate preparation system (Gasbench II) optimized for high sample throughput. Results are reported relative to the VPDB standard, and standardization was accomplished using NBS 19. The long-term precision of 13 C and 18 O values expressed as the 1-sigma standard deviation is 0:065 and 0:075‰, respectively (Spo¨tl and Vennemann, 2003). The internal structure of these speleothems was examined using standard thin-section petrography including epifluorescence microscopy.

We previously regarded the presence of speleothems as evidence of ice-free conditions above this high alpine cave, i.e. the equilibrium line altitude (ELA) was similar as today or higher. Based on age dating and stable isotope examination of several additional samples, it became evident that this assumption is too simplified. Spannagel Cave is exceptional, inasmuch as it apparently permitted local speleothem deposition also during times when most if not all of the cave system was covered by a (temperate) glacier, i.e. the ELA was at least as low as during the peak of the ‘Little Ice Age’. Two fundamentally different processes of water–rock interactions give rise to speleothem deposition in this cave system, and the stable C and O isotopes are key to identify them: (a) partitioning of soil-derived carbon dioxide into the seepage water during periods of moderately to well-developed vegetation cover, giving rise to speleothems with low 13 C values. (b) Dissolution of marble by protons released by the oxidation of disseminated sulfides (mostly in the tectonically overlying gneiss) resulting in high 13 C values in speleothems very similar to those of the hostrock. While the first mechanism is incompatible with the presence of ice above the cave, the latter process only requires the presence of liquid water in the aquifer and may continue to operate even when the cave is buried beneath a warm-based glacier. Studies of Holocene speleothems and modern dripwaters show that both processes operate hand in hand, but the pedogenic carbon dioxide source clearly being the dominant one (Spo¨tl, unpublished data). Pre-Holocene speleothem samples demonstrate that this was not the case for several older growth periods when 13 C values reflect the composition of the hostrock. These samples therefore suggest that the ground above the cave was either barren or possibly covered by ice. The only parameter that directly registers changes in atmospheric temperature above an alpine cave is the O isotopic composition.

476


Qualitatively speaking, speleothems showing high 18 O values formed from dripwaters that can be traced back to meteoric precipitation at comparably warmer temperatures than speleothems showing low 18 O values. Applying the two isotope systems in conjunction with 230Th/U dates allows us to identify the following scenarios: (a) Warm interglacial climate, high ELA, alpine vegetation above the cave: high 18 O and low 13 C values, e.g. Holocene (b) Cool interglacial or interstadial climate, ELA slightly lower than during the peak of the ‘Little Ice Age’, barren karst landscape and large parts of the cave in subglacial position: still relatively high 18 O, high 13 C values, e.g. Greenland Interstadial 14 (Spo¨tl et al., 2006). (c) Cold, stadial climate, ELA lower by several hundred meters, ice covers entire cave: low 18 O and high 13 C values, e.g. stadial preceding Greenland Interstadial 12 (Spo¨tl et al., 2006). (d) Cold, full glacial climate, ELA lower by ca. 1 km: no speleothem deposition probably because the glacier at this altitude was cold based, e.g. LGM. The predominantly interglacial records discussed in this article were obtained bothfrom flowstones, i.e. sheet-like accumulations of layered speleothems that form beneath a spatially variable dripwater source, and from stalagmites which are fed by a rather constant point source. A common feature of the former type of speleothems is the lateral variability of individual growth layers, reflecting changes in the amount of water flowing over different parts of its growing surface. As a result, the thickness of individual growth intervals as recorded by a single flowstone sample may not be representative of the overall growth rate of the flowstone. In addition, growth interruptions and related minor hiati may be the result of flow-route switching on the flowstone surface and hence

were not related to (external) hydrological changes. None of the sampled speleothems were currently active, and most of them were removed from their original growth position by former cave streams, most likely associated with flooding when the cave acted as a subglacial meltwater conduit (NB: there are no high-discharge streams in the cave today). While growth rates of speleothems from low-lying caves beneath densely vegetated areas in central Europe are typically approaching and exceeding 100 mm=yr (e.g. McDermott et al., 1999; Niggemann et al., 2003), resulting in Holocene stalagmites 50 to over 100 cm high, speleothems from Spannagel Cave are dwarfed and commonly contain multiple-growth discontinuities, similar to speleothems from high-latitude cave sites (e.g. Linge et al., 2001; Berstad et al., 2002). While local processes in the cave (e.g. spatial variability of dripwater supply) may account for minor discontinuities, prominent hiati are associated with major environmental changes outside the cave causing calcite precipitation to stop. Owing to such growth interruptions reliable depth versus age relationships are difficult to establish, a feature common in other flowstones as well (e.g. Baker et al., 1995; Wang et al., 2004). We therefore mostly restrict our discussion to individual growth intervals defined by 230Th/U dates. 31.4.1 Flowstone SPA 52 This piece of flowstone actually consists of three samples, a larger piece (SPA 52A) and two lateral equivalents of its upper part (SPA 52B, Fig. 31.4) and SPA 11, about 30 cm and 40 cm apart from the first one, respectively. Figure 31.5 shows a high-resolution stable isotope transect of SPA 52A together with the updated set of 28 230Th/U dates. According to the 230Th/U dates, about two-thirds of the samples were formed during MIS 7.3 to 7.1. The corresponding warm phase in the Alps – indicated by high 18 O values of the calcite – lasted from 211 and 189 kyr.


SPA52A

H4

477

SPA52B H4

H3

H3

H2 H2

H1

4 cm

Fig. 31.4 Slab of flowstone SPA 52A and lateral subsample SPA 52B (SPA 11 not pictured). The black lines indicate the locations of the high-resolution stable isotope traverses. Hiati are marked by H1 to H4.

The 18 O record reveals three intervals with consistently higher values separated by periods of lower values (Fig. 31.5). In one instance (100 mm above the base of the sample), this 18 O shift is associated with a macroscopic hiatus (H1 in Fig. 31.4). The periods of 18O-enriched calcite between MIS 7.3 and 7.1 can be compared to warm climate periods known from other climate archives. We chose the highly resolved alkenone record from sediment core ODP-977A retrieved from the western Mediterranean (Alboran Sea, Martrat et al., 2004). Its chronology is tied to the NorthGRIP ice core (North Greenland Ice Core Project Members, 2004). The chronology of the older section of the sediment is based on orbital tuning of the planktonic 18 O isotope curve. The lowermost 39 mm of the calcite in SPA 52A – dated at 211 to 206 kyr – records a period of rather high 18 O values depicting a slight trend towards lower values up section. The timing of this warm interval coincides with MIS 7.3, dated to be between 218 and 203 kyr in the Alboran Sea core (Fig. 31.5), which also shows an overall cooling trend. MIS 7.3 calcite precipitation abruptly ended in Spannagel Cave at 206 kyr associated

with a 2.5‰ drop in 18 O, consistent with the sharp 5 C cooling of the sea-surface temperatures (SSTs) in the Mediterranean. Despite this global cooling (MIS 7.2), calcite deposition continued in Spannagel Cave registering significantly lower oxygen isotope values. Following the sharp 2‰ rise at 65 mm above the base, warm conditions in Spannagel Cave (i.e. high 18 O values) were re-established at 195 kyr, which compares well to the MIS 7.1 signal of the Alboran Sea record ( 199 to 192 kyr). Even the (bifurcate) structure of this interglacial is similar in both records (Fig. 31.5). This period of high sea level (MIS 7.1) has been constrained to range from 202 to 190 kyr in the coastal Argentarola Cave (Italy; Bard et al., 2002) which is consistent with the chronology of SPA 52A. We note that MIS 7.1 did not reach the same high 18 O values as the preceding interglacial which may indicate an overall cooler temperature in the Alps during MIS 7.1 as compared with 7.3. By comparing sea-level data from Argentarola Cave and from a Bahamas flowstone (Li et al., 1989), Bard et al. (2002) concluded that sea level was higher during MIS 7.3 than during MIS 7.1.


478

Distance from base (mm) 0

20

40

60

80

100

120

140

160

180

200

90 H1

Age (kyr)

110

H2

H4

H3

130 150 170 190 210 1 6

0

δ13C (‰, VPDB)

4

–1

2

–2

δ13C

0

–3

–2

–4

–4

–5

–6

–6 –7

–8

–8

–10

–9 –10

–14

–11

–16

–12

–18

–13

–20

?

δ18O (‰, VPDB)

δ18O

–12

–14 –15

–22 –23

19

AI-26 (MIS 5.5)

AI-12′ (MIS 7.3)

AI-15′ (MIS 7.5)

AI-9′

AI-11′ (MIS 7.1)

15 13

AS-12′ (MIS 7.2)

11

AI-1′ AS-1′

17

AI-10′

SST (°C)

21

9 240

230

220

210

200

190

180

170

160

150

140

130

120

110

Age (ODP-977A; kyr)

Fig. 31.5 230Th/ U ages and the stable carbon and oxygen isotope record of flowstone SPA 52A as a function of distance from base. The chronology is based on a recalibration of the original 230Th/U dates of Spo¨tl et al. (2002) which resulted in a shift towards older ages (e.g., 3 kyr at 130 kyr). One data point at 148 kyr was discarded because it fell right at the major hiatus at 135 mm and most likely represents a mixture. The new high-resolution stable isotope track (150 m increments – Fig. 31.4) was micromilled a few centimetres off the location of the previous low-resolution profile. This new profile essentially confirms all major isotope features of the latter (Spo¨tl et al., 2002), but there are variations in the relative thickness of segments within this large flowstone sample, e.g. the thickness of the last interglacial calcite (see Fig. 31.4). The positions of the individual 230 Th/ U dates – shown with their 2-sigma errors – were plotted onto the new isotope profile by correlating salient features in both isotope tracks. Intervals of speleothem growth are highlighted by the horizontal grey bars. The growth phase during MIS 7 was discontinuous as indicated by the short, dashed, horizontal lines within the grey bar. Solid and vertical lines represent major growth hiati (H1–H4). The bottom curve shows alkenone-derived sea-surface temperature variations of the Alboran Sea (core ODP-977A; Martrat et al., 2004) for the period between MIS 7.5 and 5.5, and dashed lines suggest correlations between the 18 O profile of SPA 52 and the Mediterranean record. Events labelled AI and AS refer to warm and cold intervals, respectively (Martrat et al., 2004).


The remaining part of the MIS 7 section in sample SPA 52A is more difficult to interpret because of the presence of two discontinuities (H1 at 100 and H2 at 125 mm). There is clear evidence for a second cold period (reaching 18 O values lower than those during MIS 7.2) followed by a last warm interval with considerable internal isotope variability dated to 190 to 189 kyr. This period of high 18 O values may correspond to the short interstadial (AI-109 in Martrat et al., 2004) in core ODP977A immediately postdating MIS 7.1 ( 190 kyr; Fig. 31.5). Alternatively, this last warm period in Spannagel Cave could also correspond to the next interstadial (AI-99 in Martrat et al., 2004) centred at 186 kyr. In any case, no calcite was deposited subsequent to 189 kyr and during MIS 6, the penultimate glacial. Speleothem formation commenced again 53 kyr later, with a complexly structured layer of greyish calcite and high 18 O values compared to those of the warm MIS 7 periods. It follows a marked 3.5‰ decline and a recovery towards intermediate 18 O values (Fig. 31.5). This period is chronologically well bracket by 10 230Th/U dates ranging from 136 to 133 kyr. The Alboran Sea SST record exhibits a significant and prolonged interstadial between 138 and 131 kyr (AI-19) and a Younger Dryas-like cold reversal afterwards (AS-19 at 130 kyr; Fig. 31.5). This pattern of an early warming has also been reported from other archives, including the Devils Hole calcite (Winograd et al., 1997), a pollen record from Portugal (Sańchez Gon˜i et al., 2005), corals from Huon Peninsula (Esat et al., 1999) and Barbados (Gallup et al., 2002), aragonitic marine sediments off the Bahamas (Henderson and Slowey, 2000) and more recently from a speleothem record of Hulu Cave in China (Cheng et al., 2006). The marked drop in 18 O in flowstone SPA 52A is apparently slightly older than the Younger Dryas-type cold phase in core ODP-977A, which falls right within the major growth hiatus of SPA 51A at 170 mm above base.

479

The main phase of speleothem deposition during the last interglacial (MIS 5.5) lasted from 125 to 119 kyr, consistent with the highest SST values in the western Mediterranean Sea. The oxygen isotope data are the highest within the past 85 kyr in Spannagel Cave only rivalled by the earliest part of MIS 7.3 (Fig. 31.5). In contrast to the MIS 7 interglacials, the MIS 5.5 calcite is characterized by a surprisingly low variability in 18 O. The topmost calcite layer in this sample – overlying hiatus H4 – records a short interval of flowstone accretion during MIS 5.3. Figure 31.5 also shows a high-resolution 13 C record of flowstone SPA 52A. There are intervals of covariant 18 O and 13 C values alternating with segments of little or even negative correlation. We note a clear difference in the carbon isotopic composition of calcite deposited during the warm periods of MIS 7 and calcite formed during MIS 5.5. The latter depicts the lowest 13 C value (down to 10:3‰) at the time of high 18 O values (Fig. 31.5). This anticorrelation is consistent with the establishment of vegetation above this high-alpine cave site. In contrast, the covariant swings in both isotopes during MIS 7 and the generally high 13 C values suggest little if any input of carbon from the soil zone. Figure 31.6 compares the MIS 5.5 part of sample SPA 52A to that of the lateral sample SPA 52B. Both high-resolution records show remarkably stable 18 O values which decline at the end of the MIS 5.5 segment (marked by H4). In contrast, the 13 C values are highly variable and show an early trend of decline followed by a return to slightly higher values near the top (Fig. 31.6). We did not attempt matching the individual isotope peaks among the two adjacent flowstone samples. It is obvious however that the basic isotopic information provided by sample SPA 52A is representative for most parts of the ca. 4 m2 large flowstone. A lowresolution isotope profile is also available from a second lateral piece, SPA 11 (not shown graphically), and the first-order trends are essentially the same.


480

Distance from base (mm)

Age (kyr)

165

170

115 117 119 121 123 125 127 129 131 133 135

175

180

H3

185

190

195

200

H4 92.8 ± 0.5 kyr

2 0

δ13C

–2 –4 –6

–7

–8

–8

–10

–9

δ18O

–10 –11 –12

δ18O (‰, VPDB)

δ13C (‰, VPDB)

4

–2 –4 –6

δ13C

–8 –8

–10

–9 –12

δ18O

–10 –11 –12

δ18O (‰, VPDB)

δ13C (‰, VPDB)

–13

Fig. 31.6 18 O and 13 C profiles of the last interglacial portion of flowstone SPA 52. The upper two isotope tracks are a close-up of sample SPA 52A (see Fig. 31.4), while the two lower isotope traverses were analysed across the lateral equivalent SPA 52B (Fig. 31.4). The chronology is based on a recalibration of the original 230 Th/U dates by Spo¨tl et al. (2002). No 230Th/ U dates are available from SPA 52B. The stable isotope sampling interval was 150 m. Solid vertical lines represent growth discontinuities indicated by petrography and inferred from 230Th/ U dates. The top calcite layer formed significantly later than the MIS 5.5 calcite.

31.4.2 Flowstone SPA 59 This inactive flowstone piece ranges in thickness between 6 and 11 cm and was collected in the south-central part of Spannagel Cave (Fig. 31.2). The flowstone encompasses several intervals of calcite deposition separated by well developed hiati (Fig. 31.7). Only a summary of its salient features during MIS 7 and 5 is given here. Thirty-seven 230Th/U dates were determined from this flowstone which is presented in detail elsewhere (Holzka¨mper

et al., 2005). The flowstone record reveals a short growth phase between 236 and 229 kyr, with high 18 O values coinciding with the second half of MIS 7.5 as shown in the Alboran Sea SST record. The oxygen isotopic composition of this interglacial calcite is comparable to that of MIS 7.3 and 7.1 calcite of sample SPA 52A. 230Th/U ages derived from Bahamas slope sediments suggest that full interglacial conditions of MIS 7.5 lasted from 237 to 228 kyr (Robinson et al., 2002), which fits remarkably well to


MIS 5.5 MIS 7.1 MIS 7.5

2 cm

Fig. 31.7 Cross-section of flowstone SPA 59 showing the location of the stable isotope traverse (black line) and the stratigraphic position of calcite layers deposited during the penultimate and the last interglacial.

this growth phase of SPA 59. No calcite was precipitated during MIS 7.3 in sample SPA 59, which we attribute to site-specific hydrological changes (see above). The presence of a thick portion of MIS 7.3 calcite in sample SPA 52A demonstrates that this absence is indeed not palaeoclimatologically controlled. Overlying a hiatus, SPA 59 records a layer of calcite deposited during MIS 7.1 ( 192 to 199 kyr), consistent with the chronology obtained from flowstone SPA 52A (Fig. 31.5) and with a high sea level between 202 and 190 kyr (Bard et al., 2002). It is instructive to compare the isotope pattern of the MIS 7.1 calcite between the two flowstone samples SPA 59 and 52A (Holzka¨mper et al., 2005; Fig. 31.5): both 18 O records show a double peak, whereby the older peak is consistently larger than the younger one and the absolute 18 O values are comparable. The precision of the 230Th/U dates does not permit to verify or falsify the synchronicity of these isotope signals between

481

the two samples, but it is an intriguing possibility. No calcite was deposited after 192 kyr and during MIS 6, again consistent with the results from sample SPA 52A. Speleothem growth commenced again at 137 to 135 kyr, comparable to the first flowstone sample. 18 O values remained at an intermediate level nearly identical to that of sample SPA 52A ( 12‰). A thin layer of isotopically enriched calcite dated at 131–130 kyr records the onset of full interglacial conditions, and the peak 18 O value is precisely the same as those recorded in flowstone SPA 52A (and lateral equivalents, i.e. 8:5‰). Most of the MIS 5.5 calcite, however, appears to have been dissolved subsequently marked by a highly porous, residue-rich dark grey layer (Holzka¨mper et al., 2005; Fig. 31.7). It is unknown when this dissolution took place, but observations in Spannagel Cave reveal that different parts of the cave show evidence of active speleothem deposition and features of active corrosion (corroborated by hydrochemical data – Spo¨tl, unpublished data) during today’s interglacial. This situation arises from the wide range of hydrological flow systems, ranging from slow seepage flow (producing dripwaters supersaturated with respect to calcite) to fast fissure flow (permitting rapid access of rather aggressive surface waters (e.g. snow meltwater) to host rock and speleothem surfaces in the underground. 31.4.3 Stalagmites SPA 50 and SPA 51 Located only a few meters apart from flowstone SPA 52, these two stalagmites were found detached from their substrate (Figs. 31.8 and 31.10). They were probably broken off by high-discharge floods during the last glacial period but not transported more than a couple of meters as shown by their well-preserved original surface. Both stalagmites formed subsequent to the penultimate glacial and show comparable isotope records. Growth of SPA 50 started at 134 kyr (Holzka¨mper et al., 2004, 2005;


482

H3

4 cm

Fig. 31.8 Polished slab of stalagmite SPA 50. Note unconformity H3 near the base. The black line marks the micromilled stable isotope traverse to the right of a previously (manually drilled) low-resolution profile.

Fig. 31.9). In terms of both its formation period and the low 18 O values, this calcite is directly comparable to the early growth of calcite at the nearby flowstone SPA 52A (Fig. 31.5), as well as to the more distant flowstone SPA 59. Following a marked hiatus (H3; Fig. 31.8), calcite deposition recommenced at 129 kyr associated with 18 O values similar to that of flowstone SPA 52A. This MIS 5.5 calcite comprises the majority of the stalagmite and yielded a 230Th/U age of 121 kyr at the top. Because of its rather high growth rate, stalagmite SPA 50 could be examined in considerable detail (see Holzka¨mper et al., 2004). The results of stalagmite SPA 51 are reported here for the first time. This stalagmite is smaller (Fig. 31.10) and lacks the early growth phase of SPA 50. 230Th/U dates constrain its formation period from 127 to 119 kyr (Table 31.1), whereby the growth rate increased after 125-124 kyr similar to that of stalagmite SPA 50 (Fig. 31.9).

Age (kyr)

Distance from base (mm) 119 121 123 125 127 129 131 133 135 6

0

30

60

90

120

150

180

210

240

H3

δ13C (‰, VPDB)

4 2 0

δ13C

–2 –4 –6 –8

δ18O

–8 –9 –10 –11 –12 –13 –14

δ18O (‰, VPDB)

–10

Fig. 31.9 18 O and 13 C profiles of stalagmite SPA 50. The chronology is based on a recalibration of the original 230Th/ U dates (Holzka¨mper et al., 2004, 2005). The stable isotope sampling interval was 150 m. Dashed vertical lines represent changes in growth rate inferred from 230Th/U dates.


H4

483

interpretation. The latter stalagmite also records a much younger final calcite growth phase ( 56 kyr). Growth during MIS 3 is also replicated in flowstone SPA 59 (Holzka¨mper et al., 2005) as well as in two stalagmites from a neighbouring cave (Spo¨tl and Mangini, 2002; Spo¨tl et al., 2006).

31.5 PALAEOCLIMATIC DISCUSSION 31.5.1 MIS 7

2 cm

Fig. 31.10 Polished slab of stalagmite SPA 51 and location of stable isotope profile (black line). Hiatus H4 separates MIS 5.5 calcite from a thin layer of calcite formed during MIS 3. The apparent boundary ca. 5.5 cm above the base is not a hiatus.

The high-resolution stable oxygen isotope profile resembles that of its neighbour SPA 50 as well as that of flowstone SPA 52A with respect to both the low amplitude of variation and the absolute value (Fig. 31.11). The carbon isotope record of these three samples show similar overall features, i.e. 13 C values as low as 10‰ and a much higher amplitude as compared with that of the 18 O vales. The lack of an obvious match between the three carbon isotope curves suggests that this parameter reacts sensitively to minor differences between individual dripwater sites. Excursions towards high values are attributed to stochastic variability (decrease) in drip rate and progressive departure from equilibrium conditions. Parallel increases in 18 O – seen best in stalagmite SPA 51 – corroborate this

Given the fact that reliably dated palaeoclimate records from surface sediments of the penultimate interglacial in the Alps are unavailable, the flowstone records retrieved from Spannagel Cave have far-reaching significance. Flowstones are inherently more vulnerable to small changes in the seepage water supply than stalagmites. Nevertheless, both flowstones SPA 52A and 59 (Holzka¨mper et al., 2005) provide remarkable insights into the timing and progression of warm periods during MIS 7. As there is no Northern Hemisphere ice record going back to MIS 7, we chose the SST record from the Alboran Sea as a marine archive for comparison. Although this record is continuous and highly resolved, it lacks direct age control (Martrat et al., 2004). This alkenone record displays large, up to 10 C, temperature changes in the surface water of the Mediterranean Sea between warm and cold periods within MIS 7 (Martrat et al., 2004). The overall climate evolution seen in this record strongly resembles reconstructions from pollen data (Tzedakis et al., 2004), albeit at considerably higher detail. Periods during which calcite enriched in 18O was deposited in Spannagel Cave coincide with MIS 7.5 (SPA 59), 7.3 (SPA 52A) and 7.1 (SPA 52A and 59). This is also seen in the marine and terrestrial (pollen) records and provides first-time evidence of these three isotopically dated interglacial periods in the Alps. More important, the timing is consistent with other U-series

484

Lab #

2478 3253 3794 3254 3255 3829 3258 3793 3792 3830 3791

Distance from top (cm)

0.3 3.0 4.5 6.5 8.3 13.0 13.5 14.6 16.0 17.0 18.7

234 U

Concentration238U

Concentration232Th

Concentration230Th

Age

(‰)

(‰)

ðmg=gÞ

ðmg=gÞ

(ng/g)

(ng/g)

(pg/g)

(pg/g)

(kyr)

(kyr)

23.9 26.6 27.1 22.3 20.9 14.9 20.2 14.2 11.9 11.9 13.5

1.8 2.2 1.9 1.6 1.6 1.7 1.7 1.8 1.4 1.4 1.9

109.14 2.7110 3.5610 5.4210 4.0670 3.2493 3.7050 4.6840 5.5250 4.4753 6.3610

0.13 0.0027 0.0036 0.0054 0.0041 0.0032 0.0037 0.0047 0.0055 0.0045 0.0064

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