Global Palaeoclimate of the Late Cenozoic (Developments in Palaeontology and Stratigraphy)

Global Palaeoclimate of the Late Cenozoic FURTHER TITLES IN THIS SERIES 1 . A.J. Boucot EVOLUTION AND EXTINCTION RATE...

Author: V. A. Zubakov | I. I. Borzenkova

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Global Palaeoclimate of the Late Cenozoic

FURTHER TITLES IN THIS SERIES 1 . A.J. Boucot EVOLUTION AND EXTINCTION RATE CONTROLS

2. W.A. Berggren and J.A. van Couvering THE LATE NEOGENE - BIOSTRATIGRAPHY, GEOCHRONOLOGY AND PALEOCLIMATOLOGYOF THE LAST 15 MILLION YEARS IN MARINE AND CONTINENTALSEQUENCES

3. L.J. Salop PRECAMBRIANOF THE NORTHERN HEMISPHERE 4. J.L. Wray CALCAREOUS ALGAE

5. A. Hallam (Editor) PATTERNS OF EVOLUTION, AS ILLUSTRATED BY THE FOSSIL RECORD

6. F.M. Swain (Editor) STRATIGRAPHIC MICROPALEONTOLOGY OF ATLANTIC BASIN AND BORDERLANDS 7. W.C. Mahaney (Editor) QUATERNARY DATING METHODS

8 . D. Jan6ssy PLEISTOCENE VERTEBRATE FAUNAS OF HUNGARY

9. Ch. Pomerol and I. Premoli-Silva (Editors) TERMINAL EOCENE EVENTS

10. J.C. Briggs BIOGEOGRAPHY AND PLATE TECTONICS 11. T. Hanai, N. lkeya and K. lshizaki (Editors) EVOLUTIONARY BIOLOGY OF OSTRACODA. ITS FUNDAMENTALSAND APPLICATIONS

Developments in Palaeontology and Stratigraphy, 12

Global Palaeoclimate of the Late Cenozoic \(.A.Zubakov and I.I.Borzenkova State Hydrological Institute, 23 Line 2, Leningrad 199053, U.S.S.R.

ELSEVIER Amsterdam - New York - Oxford - Tokyo

1990

ELSEVIER SCIENCE PUBLISHERS B.V. Sara Burgerhartstraat 25 P.O. Box 2 1 1, 1000 AE Amsterdam, The Netherlands

Distributors for the United States and Canada: ELSEVIER SCIENCE PUBLISHING COMPANY INC. 555, Avenue of the Americas New York, NY 10010, U.S.A.

L l b r a r y o f Congress Cataloging-in-Publication

Data

Zubakov. Vsevolod Alekseevich. [Paleokllmaty pozdnego kainozora. English] G l o b a l p a l a e o c l i m a t e o f t h e late C e n o z o i c / V.A. Z u b a k o v a n d 1.1. B O r Z e n k o v a . p. c m . -- ( D e v e l o p m e n t s in p a l a e o n t o l o g y and s t r a t i g r a p h y ,

12 1 Translation o f Paleoklimaty pozdnego k a i n o z o b . Includes bibliographical references. ISBN 0-444-87309-0 ( U . S . ) 1. Paleoclimatology. 2 Geology, Stratigraphic--Cenozoic. I. B o r z e n k o v a . I. I. ( I r e n a I v a n o v n a ) 1 1 . T i t l e . 1 1 1 . S e r i e s . ac8.~4.~83513 1990 89-48 i 84 551.69--dc20 CIP

ISBN 0-444-87309-0 (Vol. 12)

0Elsevier Science Publishers B.V., 1990 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, Elsevier Science Publishers B.V./ Physical Sciences & Engineering Division, P.O. Box 330, 1000 AH Amsterdam, The Netherlands. Special regulations for readers in the USA -This publication has been registered with the Copyright Clearance Center Inc. (CCCJ, Salem, Massachusetts. Information can be obtained from the CCC about conditions under which photocopies of parts of this publication may be made in the USA. All other copyright questions, including photocopying outside of the USA, should be referred t o the copyright owner, Elsevier Science Publishers B.V., unless otherwise specified. 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. Printed in The Netherlands

CONTENTS

Preface . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

ix

Preface to the Russian edition of Pulueocl~rriatesoj’lhe Lafe Cenozoic by V.A. Zubakov and 1.1. Borzenkova (Gidrometeoizdat 1983) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

xiii

Preface to the Russian edition of The Glabul Cliimtic Events of Ihe Pleistocene, by V.A. Zubakov (Gidrometeoizdat, 1986) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

YV

.

.

Part 1 The global climatic events of the Pleistocene

..........

...........

Introduction (V.A.Z.) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Why has the interest in the past climates grown strikingly? . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . On two paradigms of palaeoclimatology . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . The main goals of this study . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

.

Section I Methodological problems of palaeoclimatology (V.A.Z.) . . . . . . . . . . . . . . . . . . . . . . .

.

.......................................... ate and palaeoclimatography . . . . . . . . . . . . . . . . 1.2. On the terms “global climatic event”, “climathem”, “climarostratigraphy” . 1.3. On the methods of high-resolution climatostratigraphic correlation a of global climatic events . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . I .4. The principles of time classification of the global climatic events: Taxonomic differences in the climato - sedimentary cycles and climathems . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.5. The two climatic regimes in the history of the Earth . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.5.1. Main features of the glacial climatic regime . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . I .5. 2. Main features of the greenhouse climatic regime . . . . . . . . . . . . . . . . . . . . . . . . . . . Resume . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

Chapter 1 The time structure of climate I .1 . On the definitions of climate, palaeo

.

Chapter 2 Deep-sea standard for global climatic events . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.1. History of climatostratigraphic study of the Pleistocene . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2 . 2 . The significance of the oxygen-isotope scale for climatostratigraphic reconstructions . . . 2.3. Systematic aspects of “ocean - continent” climatochronological correlation . The significance of geomagnetic data . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Resume . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

.

Section 11 Evidence for climatic changes in the Pleistocene .regional review (V.A.Z.) . . . . .

.

Chapter 3 Effects of global climatic events in the Mediterranean - Caspian system

.........

3.1. The Mediterranean as a new climatoparastratotype region . . . . . . . . . . . . . . . . . . . . . . . . . . 3.2. The Caspian basin as a major record of changes in humidification in interior Eurasia . . 3.3. The Azov- Black Sea basin as a standard for the climatostratigraphic sequence on the shelf off Europe ............................................................... 3.4. The Mediter ydrol of global and regional climatic cha .... Resume . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

i

3 3 5 10 13

15 15 17 18

21 22 28 32 38 39 39 45

54 65 67 69 69 76 83

93 100

vi

Chapter 4 . The loess assemblage of Eurasia as an indicator of climatic changes in the arid zone Resume . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

101 101 115 122

Chapter 5 . Middle and high latitudes of the Northern Hemisphere as a major record of continental glaciations in Pleistocene time . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.1. Russian plain . . ....................................... 5.2. Glaciated area in Western and Central Europe . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.3. West Siberia . . ....................................... 5.4. North-eastern A ... ............................... 5.5. North America . . . . ........... ........... 5.6. The Arctic and sub-Arctic . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Resume . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

125 125 132 145 156 164 172 184

4.1. The loess zone of Europe . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.2. Loess in Asia . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

. Chapter 6 . On the timing of palaeoclimates in the Pleistocene (V.A.Z.)

Section 111 The history of climate through the Pleistocene . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

...................

6 . I . Debatable problems of inter-regional climatostratigraphic correlation . . . . . . . . . . . . . . . . . 6.1.1. O n correspondence between the numbers of climathems on land and in the sea 6.1.2. O n two stratific lines in the geo-historical classification of the Pleistocene . . . . 6.1.3. Comparison of experiences in the long-distance stratigraphic correlation of the ........................ .... 6.2. Rhythm-chronological approach to the Pleistocene classification .... 6.2.1. On three types of time classification of the Pleistocene clim 6.2.2. The role of the 400 ka cycle for chronological classification of the Pleistocene

187 189 189 189 189

Resume . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

194 197 198 200 206

Chapter 7 . Climatic changes in the Early and Middle Pleistocene (V.A.Z.) . . . . . . . 7.1. Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7.2. The sixth (Ciinz) kryo-superclimathem, 1.17 - 1.0 Ma ..................... 7.3. The fifth (Giinz - Mindel) thermo-superclimathem, 1 .0 - 0.76 Ma . . . . . . . . . . . 7.4. The fourth (Mindel) kryo-superclimathem, 760- 585 ka . . . . . . . . . . . . . . . . . . . 7.5. The third “Mindel - Riss” thermo-superchathem, 585 - 350 ka . . . . . . . . . . . 7.6. The second (Riss) kryo-superclimathem, 350- 130 (170?) ka . . . . . . . . . . . . . . . Resume . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

209 209 209 210 212 213 215 217

Chapter 8 . Climatic changes in the Late Pleistocene . .... 8. I . Tyrrhenian (= Riss - Wiirn? sensu lato) megathermochron. 245 - 118 O n the time-scope of the “Riss- Wiirm” (277-244 ka) . . . . . . 8.1.1. 8.1.2. The Early Riss - Wiirm - the seventh thermo.orthoclimathem, 8.1.3. The sixth kryo.ortoclimathem, 190- 127 ka . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 8.1.4. The Late Riss-Wiirm or thermo-orthochathem 5e, 127(170?)- 117 ka . . . . . 8.2. Spatial climate reconstructions for the temperature optimum of the last thermochron (isotopic substage 5e), 125- 120 ka (I.I.B.) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 8.3. The Wiirm megakryochron, 117-15 ka (V.A.Z.) . . . . . . . . . . . . . . . . . . . . . . . . . . . 8.3.1. O n chronological models of the last glaciation . . . . . 8.3.2. The Early Wiirm, or kryo-orthoclimathem 5d.4, 117-62 ka 8.3.3. The Middle Wiirm - thermochron 3c, 62-42 ka . . 8.3.4. The Late Wiirm, 42- 13 ka . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 8.4. Spatial reconstruction of the Northern Hemisphere climate during the Late Wiirm, 20- 17 ka (I.I.B.) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Resume . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

219 219 219 22 1 225 226

Chapter 9 . Climatic changes through Late Glacial and Post.glacia1. 16-0 ka BP (I.I.B.) . . . . 9.1. Principles of the time classification of the last 16 ka . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 9.2. On the global temperature trend over the last 16 ka . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 9.2.1. Anathermal from 16 to 9 ka BP . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

251 251 256 256

227 231 231 237 240 241 246 248

vii 9.2.2. Megathermal, 9 - 5.3 ka BP . . . . . . . . . . . . . . . ..... 9.2.3. Katathermal, 5.3-0 ka BP . . . . . . . . . . . . . . . . . . . . . . . . . . . ........... 9.3. On possible causes of climate change in the Late Glacial -Holocene . . . . . . . . . . . . . . . . . 9.4. Moisture conditions in different latitude zones over the Late Glacial - Holocene: a review of empirical data . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 9.4.1. Empirical data on moisture conditions in tropical and subtropical regions between 0 and 25"N and S . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 9.4.2. Empirical data on moisture trends during the Late Glacial - Holocene between 25 and 40"N and S . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 9.4.3. Empirical data on moisture condition variations in middle (between 40-45" and igh (north of 60"N) latitudes . . . . . . . . . . . . . . . . . . . . . . . . . . . Resume . . . . . . . . . ......................................................

260 261 267

Summary (V.A.Z.) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Global climatic events - an empirical basis for high-resolution stratification . . . . . . . . . . . . . . On the causes of climatic changes in the Pleistocene . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Resume . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

291 291 298 310

215 216 286 288 294

.

Part II Pre-Pleistocene climates: Main steps of the Late Cenozoic glacial-psychrospheric regime standing . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

313

Preface . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

315

Introduction (V.A.Z.) . . . . . . . . . . . . . . . . . .

317

Chapter 10. Paleoclimates of the pre-Pliocene Cenozoic (V.A.Z.) . . . . . . . . . . . . . . . . . . . . . . . . . 10.1. The state-of-the-art of stratigraphy, geochronology and historic subdivision of the Cenozoic . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 10.2. The transition from the greenhouse - thermohaline regime to the glacial one, 48-38 Ma . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 10.3. Psychrospheric climatic regime of the Oligocene/Early Miocene, 37 - 29 Ma . . . . . . . . . . 10.4. Early-Middle Miocene optimum, 21.0- 15.3 Ma . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 10.5. Paleoclimates of the Middle- Late Miocene, 15.3 7.8 Ma . . . . . . . . . . . . . . . . . . . . . . . . . Resume . . . . . . . . . ............................................................

319

~

.

319 325 331 335 344 348

Chapter 11 Paleoclimates of the Pliocene (V.A.Z.) . . . . . . . . . . . . . . . . . . . 11.1. The Black Sea standard for the Pliocene 11.2. The Caspian Sea region . . . . . . . . . . . . . . .............................. 11.3. The Mediterranean, North-Western Eur ................... 11.4. The main steps in the Pliocene climate ................... 11.4.1. Palaeoclimates of the Early PI .................. 11.4.2. The Middle Pliocene warm cli ................... 11.4.3. Palaeoclimates of the Late Pliocene (Villafranchian) 3.65 - 1.0 Ma . . . . . . . . . . . 11.5. Tentative reconstruction o f climatic conditions for the Northern Hemisphere during the Middle Pliocene (I.I.B.) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . ResumC . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

351 35 1 363 314 384 386 388 391

Summary (V.A.Z.) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

401

396 399

.................................................................

403

References to Part i . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

405

References to Part I i . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

443

Subject index . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

455

Acknowledgements

This Page Intentionally Left Blank

PREFACE The history of this book goes as far back as 1979, when Professor M.I. Budyko, Chairman of the Climatic Changes Research Department of the State Hydrological Institute, suggested that the authors should try to develop reconstructions of climatic environments of the Late Cenozoic optima. Budyko thought that such reconstructions could be used as a kind of “palaeoanalogue” in forecasting the 21st century climate. When starting work on this subject, the authors had not immediately realized what a mountain of unsolved problems they would need to overcome along the way. Some problems lying on the surface were associated with the almost undeveloped methodology for quantitative calculations of the parameters of the past climate. Other more latent, and therefore especially puzzling, problems ran into the absence of a reliable chronology of climatic changes and, hence, ambiguous correlation of “climatic signals’’ from place to place. The authors considered the latter to be a matter of priority because they understood that palaeoclimatic correlation that does not rely on the valid chronology would suffer the fate of constructions set up without foundations. Thus, the primary goal of our investigations has been directed at a regional systematization of the available information on the Late Cenozoic climatic changes and their classification in time. In 1980- 1982 we were engaged in preparing the first summary Palaeoclimates of the Late Cenozoic (published in 1983 by Hydrometeoizdat) presenting a preliminary review of the state-of-the-art in this field. In 1983 - 1985 the book The Globd Climatic Events of the Pleistocene was written, being published at the end of 1986. This study treated more thoroughly the data covering the time interval from 1 .O Ma to 10 ka. The English version combines the material of both Russian editions, but is organized differently, and is published as one volume. Part I includes general methodological issues and comprehensive characteristics of the Pleistocene climate. The climatic history of the Pliocene and the Late Miocene (15 - 1.1 Ma) is given in Part 11. The main content of the book consists of regional essays on climatostratigraphic sequences in 12 regions of the Northern Hemisphere. Most elaborately treated are the regions in the USSR territory (such are 8 out of the 12), which are covered by the authors’ personal information. In composing the essays on the regions outside this country, the authors relied on the literature found in the All-Union Geological Library, Leningrad Branch. Although the funds of this library are good, they could not include all the latest palaeoclimatic information from abroad. Therefore, there sometimes was the inevitable informative hiatus in the book, which sometimes has been filled with information from reprints and books the authors received directly from their foreign colleagues. Taking this opportunity the authors express their sincere gratitude to many foreign scientists for the latest material they were sent,

X

among them Drs. E. Sundquist, H . Flohn, V. Sibrava, A. Berger, E.M. Van Zinderen Bakker, H. Faure, J . Chaline, D. Dreimanis, L. Lindner, N. Shackleton, G. Smith, M. Cita, A. Azzaroli, W. Dansgaard, Liu Tung-sheng, P . Ciesielski and J. Kutzbach. It sounds paradoxical but one of the difficulties the authors came across in writing the book has been both the excessively immense bulk of palaeoclimatic information scattered in the ocean of papers, books and collective monographs, and the inadequacy of every single informative source. This has severely impeded the treatment of information and made quotation particularly difficult. While working on the book the authors have used about 4,000 sources. For obvious reasons they could not however list more than 1,000 items as references chosen, alas, quite arbitrarily, for which they hope to be excused by their colleagues and readers. Trying to present as much information as possible within the confined limits of the book, the authors have compressed it by constructing rather complicated climatostratigraphic tables summarizing the data on whole regions and subregions. Judging by the local geographical names mentioned in tables, an experienced reader can always decide whether the intermediate correlation of data has been done properly. The same principle of the greatest possible generalization has served for selecting the figures for the book. The majority of them are original. The figures derived from the literature have been redrawn to a certain extent, either by being supplemented or, on the contrary, simplified. When composing the book, the authors, one of whom is a specialist in stratigraphy and the other a climatologist, partitioned the book according to their interests. Sections 8.2, 8.4 and 9, concerned mostly with climatological aspects, have been written by 1.1. Borzenkova, and the remaining part by V.A. Zubakov. It stands to reason that all general ideas and principal issues of the book have been discussed and agreed upon by both authors, but at the same time this does not exclude some differences in the authors’ views on particular problems. For example, there is a certain disagreement between the authors as to the role of vulcanicity and solar -Earth relations in climatic change. As to interpreting the history of the past climate, i.e. a methodological approach to understanding climatic records, the reader evidently will immediately notice that the authors of the book adhere to a philosophy unusual for the English-speaking scientists. Most of the scientists outside this country, particularly in America, are followers of a pragmatic chronostratigraphic concept, which has been recommended for general use by the International Stratigraphic Guide (Hedberg, 1976). The overwhelming majority of Soviet researchers in the past and at present rely on the “event” concept, when interpreting the history of the Earth, which corresponds better t o dialectical methods. Climatostratigraphy, the theoretical foundations of which have been discussed in this book (in Part 1 and in even more detail in Part 2) expresses most explicitly and probably most fruitfully the “event” concept. Climatostratigraphy is incompatible with chronostratigraphy and attracts different opinions. For instance, the North American Code (1985) denies in principle the existence of climatostratigraphy. Chronostratigraphy is however sometimes addressed in similar terms. From our point of view, sound ideas have appeared in the book Catastrophes and Earth

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History: A New Uniformitarianism edited by Berggren and Van Couvering, which has been translated into Russian (Mir, 1986). One of the authors, D. Ager is in general inclined to renounce chronostratigraphy, which he justly calls an ugly hybrid of time and rocks, which only introduces confusion into the entire system and the discussion as a whole. We think that the problem of the place and role of climatostratigraphy should inevitably become central in the methodological discussion of the “event” and chronostratigraphic approach to the Earth’s history. It is quite clear that the development of climatostratigraphy is unavoidably associated with some innovations, including the introduction of new notions and terminology. The reader will come across some of them in this book. Using the subject index, it is possible to find the definitions of them in the text. The most frequently used terms are thermomer and kryomer, which were suggested by Liittig (1954). Following his example, the authors write this term with the first letter “k” as “kryomer” (from the Greek root “kryos”) instead of “cryomer”. We take this opportunity to express our sincere gratitude to Dr. E . Sundquist and Prof. D. Ager for the idea of translating this book into English and to Elsevier Science Publishers for publishing the translation. The book has been translated by Ms. S.F. Lemeshko (Preface, Introduction, Chapters 1.1, 1.2, 1.3, 4, 6, 7, 8.1, 8.3), Ms. A.Ya. Minevich (Chapters 1.4, 5.3, 5.6, Summary of Part I), Ms. R.V. Fursenko and Ms. R.E. Sorkina (Chapters 2 , 3, 5.1, 5.2, 5.4) and Ms. V.G. Yanuta (Chapters 8.2, 8.4, 9, References). The authors are grateful to the translaters as well as to Ms. M. Kalesnik and Ms. 0. Dyagileva for their assistance in preparing the book. Vsevolod A. Zubakov Irena I . Borzenkova


PREFACE TO THE RUSSIAN EDITION Palaeochnates of the Late Cenozoic by V.A. Zubakov and 1.1. Borzenkova (Gidrometeoizdat, 1983)

This study is aimed at developing a climatochronologic basis of climatic reconstructions of the recent geological past within the climate monitoring programme. The book summarizes extensive and varied interdisciplinary material with primary palaeoclimatic information, most of which has been published since 1977, since information before that time is almost inaccessible for the broad public. The subject of the study is the surface of the entire planet; to be more accurate, the sections of the Late Cenozoic loose sedimentary cover, studied most thoroughly from its stratigraphic and chronologic points of view. In selecting the historical climatic (palaeoclimatic) series, preference has been given to highly informative oceanic and continental sections in the USSR territory, particularly in the areas adjacent to the Black Sea, the Sea of Azov and the Caspian Sea as well as in the areas of West Siberia, which have long been studied by one of the authors of the book. In describing the Pliocene and Pleistocene palaeoclimates, particular at tention has been paid to the Black Sea areas, which can be taken as a standard not only for the USSR but possibly for the entire Eurasia. The Black Sea sections serve most conveniently for establishing time relations among the Mediterranean - oceanic, pluvial Caspian Sea and glacial - periglacial north-European events. In defining the Late Pleistocene and Holocene climates, however, the greatest attention has been paid to the latest data covering territories outside the USSR, including those in Africa, Australia and America. As a result of the work, a preliminary global climatic chronologic scale for the last 7 Ma of the Earth’s history has been developed. This scale forms a basis of detailed palaeoclimatic generalizations, in particular of developing analogous palaeohistorical models of the future climate. Some of the principles governing changes in the global climate of the Late Cenozoic have also been found. The authors express their gratitude to Prof. O.A. Drozdov for the useful discussion of the book, to L.N. Mikhailova for the assistance in preparing the manuscript, to Ye.N. Ananova, N.V. Bogatina, V.K. Vlasov, N.S. Volkova, L.A. Dorofeeva, O.A. Kulikov, M.V. Muratova, V.I. Pavlovsky, S.A. Pisarevsky, V.I. Remizovsky, Ya.1. Starobogatov, G.I. Hutt for processing material obtained during field work. The authors acknowledge highly the attention paid to this study by Prof. M.I. Budyko, Corresponding Member of the USSR Academy of Sciences and Prof. G.G. Martinson, who read the manuscript and made valuable comments.

Leningrad, March 1983


PREFACE TO THE RUSSIAN EDITION The Global Climatic Events of the Pleistocene by V.A. Zubakov (Gidrometeoizdat, 1986)

Not only scientists but governments as well are nowadays anxious to know what kind of climate there will be in the 21st century. This question, however, cannot be answered without the knowledge of the causes of natural climatic changes in the past. Therefore, at present the study of the history of climate turns out to be specially addressed to a wide circle of specialists. Climatologists and oceanologists, geologists and biologists, chemists and physicists from different countries work hand-in-hand to solve this problem. An important part of these interdisciplinary studies is to outline the time structure of climate or in other words to find principles governing climatic change in time, which could be used in forecasting climate evolution. It is quite apparent that such work should be started with the Pleistocene, stretching over the last million years of geologic history, which are covered most extensively and reliably by palaeoclimatic information. This book is a continuation and further development of the study Palaeoclimates of the Late Cenozoic by Zubakov and Borzenkova, which was published by Hydrometeoizdat in 1983. Because of chronological limits, the problem has been treated in the present book more thoroughly. The conclusion made earlier has been checked on the basis of the latest information published in 1981 - 1985. Great attention has been paid to methodological problems, such as classification and rhythms of climatic events as well as the climate genesis. A valuable contribution to the work has been made by 1.1. Borzenkova, who wrote sections 8.2, 8.4 and 9.4 devoted to spatial analysis of the climate during two temperature optima (the Late Riss - Wurm and the Late Atlantic) and to the time when the Late Wurm glaciation reached its climax. In preparing the book about 2,000 literature sources have been used. Since the reference list cannot include all of them, some references (the name of the scientist and the year of publication) are given in the text itself. The authors use this opportunity to express their sincere gratitude to Prof. M.I. Budyko for supporting the study and to Ms. M.N. Kalesnik for the assistance in preparing the manuscript. Prof. O.A. Drozdov and Dr. V.D. Dibner took the trouble to read the manuscript and made many valuable and useful comments, which with gratitude were taken into account in editing the book.

Leningrad, February I984


PART I

THE GLOBAL CLIMATIC EVENTS OF THE PLEISTOCENE


INTRODUCTION

Why has the interest in the past climates grown strikingly? As early as the beginning of the 20th century Arrhenius (1903) an1 Chamberlin (18993 paid attention to the probable warming of the climate in response to CO, release into the atmosphere in the process of fuel burning. It was, however, only in the 1970s, when the rate of CO, increase reached 18 x 1015 g/yr, that the problem of man-made pollution of the atmosphere and its climatic effects became a subject of great concern. This was in particular encouraged by the works of Budyko (1972, 1974) as well as Manabe and Wetherald (1980). The important landmark in this field turned out to be a series of international meetings of experts that took place in Stockholm in 1971, in Geneva in 1979 (World Conference . . ., 1979), in Leningrad in 1981 and 1983 (Anthropogenic Climatic Change . . . , 1984), in Florida in 1984 (Sundquist and Broecker, 1985) and the same year in the German Democratic Republic (Kondratiev, 1985) and in Villach in 1985 (An Assessm e n t . . ., 1985). At all these authoritative forums, and particularly at the joint conference of ICSU/UNEP/WMO in Villach, it has been emphasized that climatic warming is inevitable due to increasing atmospheric concentrations of CO, and other greenhouse gases (nitrogen oxides, methane, ozone, freons and water vapour). Over the last hundred years the CO, concentration in the atmosphere has grown from 270 ppm to 340 ppm (Oeschger et al., 1985). I t is expected that by the year 2030 the combined concentration of all greenhouse gases will reach a value equivalent to a doubled CO, concentration compared with the pre-industrial level. As a result, the global mean air temperature will rise by 3 f 1.5"C. In high latitudes, the future warming will, however, be two to three times greater. Dramatic changes are expected in atmospheric precipitation and river run-off: these will decrease in steppe areas and increase in high and tropical latitudes. The ocean level must also rise by about 0.7 t 0.5 m. Mankind has not endured such climatic perturbation for at least the last four thousand years. Therefore, the conference in Villach, outlining great uncertainties in predicting the global and regional distributions of precipitation and temperature, recommends that governmental and financial organizations actively support the study and interpretation of the past climatic and environmental history with due regard to the climate - atmosphere - ecosystems interactions ( A n Assessmen? . . ., 1985). There are three independent approaches in estimating the possible climatic changes produced by sharply increasing carbon dioxide concentrations (Fig. I. 1). The first approach proceeds from the analysis of the data from instrumental observations for the last 100 years. These served to obtain the relationships between changes in the global thermal regime and the spatial distribution of temperature and

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Forecast of energy development and population g r o w t h

Anthropogenic climate change

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FUTURE CLIMATE FORECAST

analog :limate

Knowledge and prediction of natural climatic variations

models

The use of empirical palaeoclim a t i c reconstruct ions with n-fold C02 content as a probable f u t u r e

Palaeoclimatology

Fig. 1.1. Three approaches to the estimation of possible climatic changes in the 2lst century.

atmospheric precipitation (Vinnikov and Groisman, 1979; Vinnikov and Kovyneva, 1983; Kovyneva, 1984). The second approach suggests the use of different climatic models, which, given certain carbon dioxide concentrations, allow us to obtain both the thermal and moisture conditions in different regions of the Earth. This approach has mostly been developed by American scientists (Manabe and Bryan, 1985; Manabe and Wetherald, 1980; Atmospheric carbon . . ., 1985; Detecting . . ., 1985). The third approach applies empirical reconstructions of the past climate to those intervals of the Late Cenozoic, when atmospheric carbon dioxide concentrations were similar to those expected in the 21st century. The development of this approach was started by Budyko, who used Sinitsyn’s palaeoclimatic maps (1965) for the Eurasian Pliocene as a palaeoanalogue for constructing a prognostic map of the USSR for the mid-2lst century (Budyko et al., 1978). The Soviet national programme for studying climate devotes particular attention to the palaeoclimatic approach. The State Climatic Monitoring System includes gathering information on climatic fluctuations and changes “over the last hundreds and thousands of years and more remote time intervals” (Izrael, 1984). The State Committee for Hydrology and Meteorology and the Academy of Sciences of the USSR every year holds a symposium and publishes numerous information devoted to this subject. The collected data cover not only the territory of this country or of the Northern Hemisphere but also the Antarctic, this acknowledged “boiler” of the Cenozoic climate, as well as the lower latitudes, particularly the so-called energyactive oceanic zones (Marchuk et al., 1981) that are found at the interface of the warm and cold waters. Palaeoclimatic studies devote an ever greater attention to

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Fig. 1.2. The principle of application of past climate reconstructions for time intervals with inferred atmospheric CO, concentration as “palaeoanalogues” of the 21st century climate (after Budyko, 1982).

cartographic reconstructions of warmer climates, especially by using narrow time slices of the past temperature optima that can yield factual information on geographical distribution of temperature and atmospheric precipitation independent of model calculations. Budyko (1984) suggested using such reconstructions as specific “palaeoanalogues” for the future climatic conditions (Fig. 1.2), i f , of course, their different versions agree with each other and with model results. It is natural that to present a reliable palaeoclimatic reconstruction is possible only if a sequence of climatic events has already been chronologically and hence stratigraphically corroborated and their regional sequences have been correlated globally. Therefore, the appearance of prognostic climatology turned out to be a great impetus for the theoretical improvement of palaeoclimatology and the development of its climatostratigraphic basis.

On two paradigms of palaeoclimatology

Palaeoclimatology as a division of historical geology has already been developing for 150 years. Until now its development has been stimulated exclusively by the demands of geologic survey, since the knowledge of the past climates is absolutely indispensable in revealing the principles governing the processes of weathering and sedimentation. The theoretica1 grounds of palaeoclirnatology as they have been formed by the mid-20th century, can be expressed by a paradigm, whose sense can be rendered by the following concise points: (1) the principle source of palaeoclimatic information is the continental data; (2) the palaeoclimatic parameters are estimated proceeding from the actualistic principle (uniformitarianism);

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(3) the time sequence of palaeoclimates is determined by the available biochronological scale; (4) the global palaeoclimate is the average climate of a certain subdivision of the general chronostratigraphic scale (stage, series, system); (5) the past climate was very stable; its changes were slow and occurred at a rate that did not exceed that of mountain growth and of the accompanying processes (volcanism, changes in the configuration of the continents, seas and sea currents). An exception was the Pleistocene climate that changed rapidly and from a geologic point of view was “abnormal” because of the extremely high hypsometric position of the continents that had never been observed before the Pleistocene. The scientific and technological revolution in the field of the Earth sciences, which started in the 1960s, has fundamentally changed the position of palaeoclimatology . A subdivision of historical geology, which used singularly descriptive qualitative methods, today it has rapidly turned into an independent synthesizing scientific discipline of great practical value, which has developed its own quantitative methods. It is important to mention that about 90% of the available palaeoclimatic information has been obtained during the last 10 to 15 years as a result of numerous national and international projects such as DSDP, CENOP, CLIMAP, PIGAP as well as the projects carried out within the framework of the International Geological Correlation Programme (IGCP), in which large scientific collective bodies take part. Suffice it to say that DSDP alone, which covers the entire World Ocean, has published more than 70 volumes of Initial Reports, each containing from 400 to 1,000 pages of texts and tables. A general study evidently cannot succeed in summarizing the ever-increasing bulk of factual information, even though more than 100 papers concerned with palaeoclimatology appear every year. This “explosion” of palaeoclimatic information has occurred as a result of introducing into practice new techniques for drilling and deep core sensing as well as the replacement of a “fixist” geological concept by a “mobilist” one. The problems of palaeoclimatology are treated quite differently in recent monographic studies (Monin and Shishkov, 1979; Frakes, 1979; Schopf, 1980; Lisitsyn, 1980; Budyko, 1980, 1984; Seibold and Berger, 1982; Ushakov and Yasamanov, 1984; Yasamanov, 1985; Berger and Crowell, 1985; Sundquist and Broecker, 1985; Hecht, 1985) as compared with the studies of the 1930s - 1970s, which used only continental data. Therefore, we can say that the boundary of the 1960s - 1970s saw the rise of a new science of palaeoclimatology, which was concerned with studying the global principles of the evolution of climate, treating them like interactions among the atmosphere, hydrosphere, kryosphere and lithosphere as they are, first of all, recorded in deep-sea sediments. There is, however, another tendency in the development of palaeoclimatology, i.e. its differentiation into geological and prognostic branches. The first is a traditional branch, whose development has been determined by the demands of geological practice. The study of the past climates is absolutely necessary in revealing the principles that govern the processes of weathering and sedimentation and, consequently, of migration and accumulation of chemical elements. Palaeoclimatic analysis is necessary for forecasting the survey of mineral deposits and placers of

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noble and rare metals. At the same time geological palaeoclimatology is quite satisfied with the so-called “average climate” corresponding to time intervals of the standard stratigraphic scale or even to the local lithostratigraphic units. It is in essence the reconstruction of the average climate of the units of the international stratigraphic scale or local stratigraphic schemes that is the purpose of geological palaeoclimatology. The number of the units and the accuracy of biostratigraphic correlations limited by the time extent of the stratigraphic and biological zones (from 0.3 - 0.5 Ma to 1.O - 1.5 Ma) appear to be the natural boundary that also limits the palaeoclimatic reconstructions. In practice, it is impossible to detect by traditional methods such changes in global climate as those occurring over a time interval of less than one or two million years. It is possible to identify any number of short-term climatic variations in such concrete carefully studied sections as the DSDP site 289 (Woodruff et al., 1981) or the Kastel section of the Permian formation (Andersen, 1982), but it is practically inconceivable to correlate them globally within a framework of the traditional geological palaeoclimatology. Meanwhile palaeoclimatology has been confronted with new problems put forward by life itself, i.e. with the necessity of forecasting climatic changes produced by man’s economic activities. That has given rise to a “prognostic” branch of palaeoclimatology. The new situation has altered all the emphasis in studying the past climates. Thus, geological palaeoclimatology is mostly interested in the remote epochs, when sedimentary rocks were forming, while the “prognostic” palaeoclimatology is more concerned with the geological past of the last thousands, tens of thousands and millions of years. In order to forecast the tendency of climate development it is necessary to disclose the principles governing the natural climatic fluctuations in time, their causes and mechanisms. Therefore, prognostic palaeoclimatology is aimed not so much at reconstructing climatic conditions averaged over long time intervals as at determining the dynamics of certain climatic fluctuations, revealing extreme climatic situations and finding causal - temporal relations in interactions of the atmosphere, hydrosphere, kryosphere, biosphere, lithosphere and cosmic space. It is an acute problem to develop thoroughly the history of climate, particularly of the climate of recent geological time intervals, namely the Holocene, Pleistocene and Pliocene. It would have been better to call the prognostic palaeoclimatology ‘‘historical climatology” if this term had not already been applied to the studies devoted to the climate of the last two or three thousand years. It is natural that for the development of prognostic palaeoclimatology it is necessary t o have a new research apparatus, more accurate than that of the geological palaeoclimatology , to introduce new scientific ideas and research techniques as well as to have close contacts with adjacent scientific disciplines. This process of theoretical reformation accompanied by the advancement of new ideas and hypotheses as well as new methodological principles (model calculations, factor analysis, etc.) and even philosophical premises has stimulated the appearance of an entirely new paradigm. In the author’s opinion, its sense can be expressed as the following: (1) It should be acknowledged that the basic source of palaeoclimatology information is the deep-sea data, particularly those obtained by unconventional

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statistical methods such as the isotope and microfaunal ones; the continental data should be associated with the oceanic standards. (2) The history of climate is the alternation of relatively long periods of stable climatic conditions and brief periods of profound and fast climatic changes, which different authors call either climatic steps or climatic crises, etc. It is evident that the points where temperature changes its trend are maximally isochronic and the interconnecting lines can be considered to be the global climatic “signals”. The global climatic events of different duration (climatochrons) separated by these points form a more rational basis for palaeoclimatic reconstructions as compared to the “average climate” of the biochronologic units. ( 3 ) The classification of the past climates should be original, natural and detailed.

Fig. 1.3. Location of the main key sections (points) and groups of sections (squares) used in the review: I - Terra Amata, Grotte Lazarett, Grotte Vallonnet, Grotte Rafael, Grotte Prince; 2 - Latium Tiber River; 3 - Vrica, Santa Maria di Catanzaro, Le Castella; 4 .- Mallorca Island; 5 - TenaghiElrigen-Geroevskoc, Chokrak Phillipon; 6 - cores: RC 9 - 181, KS 09, TR 171 - 127, Alb-189; 7 Lake, Cape Tuzla, Malyi Kut Bay, Liman Trokur, Cape Krotkov; 8 - Tsvermagala Mountain, Platovo, Shirokino, Nogaisk, Port Katon, Margaritovka; 10 Duzdag, Tskhalminda, Ureki; 9 Karadzha, Baku, Zykh, Azykh; I 1 Manych Strait; 12 - Cherny Yar, Kopanovka, Mechetka, ZaDniester myany, Tsaganaman; 13 - Zavadovka, Kaydak, Kanev, Priluki, Lubny, Tiligul River; 14 River; Tiraspol, Molodovo, Korman, Khadzhibei; I5 - Paks; 16 - Gunz, Mindel, Riss and Wurm Fergana; 19 - Ordoss Plateau: Luochuan; 20 Rivers, Mauer; 17 - Tadjik Depression; 18 Choukoutian; 21 - Byelorussia - Lithuanian Region: Shklov. Aleksandriya, Korchevo, Mosty, Zhidin; Cheremoshnik; 25 Mga River, 22 - Akulovo, Chekalin, Bryankovo; 24 - Rostov (Nero Lake) Grazhdansky Prospect, Leningrad; 26 - Tegelen, Bavel, Leerdam, Eem River, Amersfoort; 27 Hamburg - Schleswig Holstein - Skaerumhede; 28 - Ehringsdorf, Taubach, Bilzingsleben; 29 Grande Pile; 30 Hokne, Clacton on Sea, West Runton, Bobbitshole, Wolston; 31 - Kozi-Grzbiet, Barkowice-Mokre, Ferdynandow, Bloni; 32 - Grabowka, Tychnowy, Grudziadz; 33 - Altai Steppe Plateau; 34 - Irtysh River, Samarovo, Semeika; 35 Belogorie Hill; 36 - Salekhard, Salemal, PyakYakha, Shchuchya River; 37 - Krasnoyarsk, 38 - Vorogovo, Panteleev Yar, Oplyvny Yar, Zavalny Yar, Podkamennaya Tungusska; 39 - Mirnoe, Bakhta River, Pupkovo; 40 - Berelyekh River, Kutuyakh, Bolshaya Chukochiya, Malyk -Sienskaya Depression; 41 - Avlekit - 0 k h o t a River; 42 Malyi Anyui Rivers; 43 - Ayon Island, Valkarai River; 44 - Krest Bay, Yanrakinot; 45 - Karaginsk Kotzebue Bay; 48 - Igarka, Ermakovo; 49 - Ust Island; 46 - Oiyagoss Yar, 47 - Nome River Yenisei Port, Kazantsevo, Karginsky Mys, Malaya Kheta; 50 - Yavai Peninsula; 51 - absent on scheme; 52 - Oktyabrskoi Revolutsii Island; 53 - Vastyansky Kon’, Markhida; 54 KipievoRodionovo; 55 - Ponoi River, Chapoma River; 56 - Wedel Yarlsberg Land; 57 - Fjasanger-Karmou; Clyde Foreland, Baffin 58 - Tingstade Trask Lake, Gotland; 59 - Camp Century; 60 - Dye 3; 61 Medicine Hat, Welsh-Valley, Wascana Creek; 64 Island; 62 - Morgan Bluffs, Banks Island; 63 Great Lakes Region, Port Talbot, Cherrytree, Plum Point, etc.; 65 - Marine terraces: Princess Anne, Socastee, Canepatch-Talbot, Waccamow; 66 - Nebraska Iowa Region: Hummel Park, Cedar Bluffs, Nickerson, Afton, Hartford, David City, etc; 67 - Great Basin Lakes (Bonneville, Lahontan, etc.); 68 - Searles Lake; 69 - St. Clemente N a n d ; 70 Cook Inlet; 71 Bermuda Island; 72 - Chad Lake; Lakes Alerce and 73 - Omo River - Turkana Lake - Olduvai Gorge; 74 - Lancaster Lake; 75 Taiquemo, Chile; 76 - Argentina Lake; 77 - Fuquene - Palasio Lake; 78 - Lakes Torrance, Frome and Leake; 79 - Euramoo and Lynch Craters. Deep sea sites and cores: 80 - V 28 - 238; 81 - V 28 - 239; 82 - P 6304 - 9; 83 - V 16 - 205; 84 - V 22- 174; 85 - M 13519; 86 - V 23- 100; 87 - DSDP 397 and M 12309; 88 - V 30-97; 89 V 29- 179; 90 - DSDP 552A; 91 - DSDP 5028; 92 - DSDP 503 and RC 10- 65; 93 - DSDP 504; 94 - V 19-30 and DSDP 157; 95 - DSDP 357 and 517; 96 - RC 1 1 - 120; 97 - M 984; 98 - RC 13229; 99 - DSDP 47; 100 - DSDP 55; 101 - DSDP 167; 102 - DSDP 277 and 279; 103 - DSDP 281. -

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To satisfy these requirements it should have a factual climatostratigraphical and climatochronological basis independent of the biostratigraphical one. The basic unit of such classification is a climathem. (4) A high-resolution technique of timing climatic events and signals and their comparison with other geological (volcanism, etc.) and astronomical events form an empirical approach to revealing the causes and mechanisms of climatic changes. The model calculations and scenarios should serve to check and control the empirical conclusion. The causes of climatic change can be considered reliably established only if the results of two independent approaches have coincided. (5) The processes of the oceanic circulation determining redistribution of carbon dioxide and water between the atmosphere and hydrosphere are most important in the mechanism of any climatic changes whatever their duration. The rhythms of these processes are irregular because they are triggered by external periodic forcings associated with the Earth’s rotation and with changes of orbital parameters within the Solar System and the galactic system. The shortest climatic - oceanic rhythms are those of El Niiio that last several years to tens of years, whereas the longest rhythms are the alternation of the so-called climatic - oceanic regimes (the “greenhouse” or thermohaline and glacial or “psychrospheric” ones), the duration of which is about 50 Ma to 250 Ma. The palaeoclimates of the Cenozoic represent different versions of these regimes. (6) The principles of uniformitarianism should be used very carefully in elucidating the history of the past climate, particularly in regard to the “greenhouse” climatic - oceanic regime. It is quite probable that the principle of historical comparison is more useful in understanding the history of the climatic past, because it combines the elements of both actualism and catastrophism. Not a single point of the above-mentioned problems is yet reliably corroborated or, moreover, generally accepted. However, all of them have already been discussed more or less thoroughly in the specialist literature of recent years. The authors have combined them into a single system, believing that even such an intuitive description of separate elements of the new concept (paradigm) can stimulate necessary discussion.

The main goals of this study The main goal set by the authors is the systematization of the available information on the climatic changes in the Late Cenozoic on the basis of climatostratigraphic classification of sediments and subsequent chronological classification of different climatic events. An attempt at doing this should naturally be started with the data covering the Pleistocene, since the Quaternary geology is a cradle of climatostratigraphy and at present the Pleistocene is still its main object. The book considers 12 regions of the Northern Hemisphere continents, the climatostratigraphic sequence of which has been studied more thoroughly. All regional essays are accompanied by general climatostratigraphic schemes that are as far as it is possible associated with the deep-sea standard, i.e. oxygen-isotope stages. It is

believed that the considered factual information (Fig. 1.3), partly covering also the Southern Hemisphere (Chapters 7, 8 and 9), will make it possible to carry out the global climatostratigraphic correlation and present in the first approximation a draft version of the general climatochronologic scale of the Pleistocene. The second goal of the study is an attempt at making rough climatic reconstructions for some warm intervals of the Late Cenozoic, i.e. the thermal optima of the Holocene and the Riss- Wurm (Chapters 8 and 9). It is quite obvious that such reconstructions can properly be made only by collective scientific bodies. Therefore, the authors consider the present reconstructions only as the very first step in advancing in this direction. And finally, the third goal is a preliminary discussion of the possible causes of climatic change. Since the most complete and reliable information on climatic fluctuations is available on the Holocene and the Late Glacial, this task mostly concerns the last twenty-thousand-year interval (Chapter 9 and Summary).


SECTION I

METHODOLOGICAL PROBLEMS OF PALAEOCLIMATOLOGY


Chapter I THE TIME STRUCTURE OF CLIMATE 1.1. On the definitions of climate, palaeoclimate and palaeoclimatography Climatology understands climate as a combination of different types of weather, i.e. a certain recurrence of temperature, precipitation, cloudiness, wind, river runoff and so on, that is averaged statistically for a certain locality (the local climate) or globally over a time period of at least thirty years (Inadvertent . . ., 1971). This definition, however, cannot be applied to the past climate. Therefore, the most general definition of the climate (the global one) that can be applied to the present and past climates might be the following: the climate is a state of the atmosphere - hydrosphere - kryosphere system and its fluctuations in time that are represented by the redistribution of heat and moisture on the Earth’s surface. Functionally, according to information sources, in such a broad sense the climate can be divided into the modern, historical and geological climate (Table 1.1). The subject of the present study is the geological climate, which is usually called the “palaeoclimate”. The palaeoclimate can be divided into the local and the global. The local climate is described by the climatic parameters that are reconstructed from the sediments of the local (regional) stratigraphic units: formation ( = suite), layer, member, biozone or horizon, i.e. the seasonal temperature fluctuations (summer and winter temperatures), atmospheric precipitation and its seasonal patterns, prevailing winds, river run-off, water temperature, sea currents and so on. The global climate is described by other parameters, including first of all the latitudinal climatic zonation (the temperature gradient between the equator and high latitudes), the extent of continentality, the atmospheric and oceanic circulation and so on. These parameters are obtained by summing up and averaging the data for the field of sediments belonging to a certain correlatable global unit of the general (international) stratigraphic scale, i.e. a system, series or stage. Both the local and global palaeoclimates are the “average climate” of a certain complex of layers, whose age varies from place to place. For instance, at one locality climatic information refers to the lower portion of the formation or stage, at

Table 1. I . Climate and its functional definitions ~ ~ _ _ _ _ ~ __ ~Climate, in the broad sense, is a state of the atmosphere - hydrosphere - kryosphere system and its temporal changes:

( I ) Modern climate is a statistically averaged sequence of different types of weather as they are recorded in the data sets of instrumental observations; (2) Historical climate is a sequence of different type5 of weather, including extreme fluctuations a \ they are recorded in chronicals and archaeologic evidence; (3) Geologic climate or palaeoclimate is a state of the climatic system derived from geological data. -

-

_.

16

another to its middle portion, at a third to its upper portion, whereas at a fourth locality its stratigraphic position relative to the boundaries of the unit can be vague, The statistical information collected in this way is often random and the “average climate” turns out to be non-representative. Fig. 1.1 shows a unit of palaeoclimatology aimed at restoration of local palaeoclimates and local palaeoclimatic events, which is conditionally called “clirnatogruphy”. It is first of all associated with the development of a methodology for calculating palaeotemperature and other palaeoclimatic parameters using lithofacies, palaeontological - ecological and geochemical - isotope data, with the help of which the actualistic principles revealed in other geographical disciplines are usually extrapolated into the past. The methodology for such a reconstruction has been developed for a long time and many publications have been devoted to this subject, which are impossible to list here. Now we shall note that geological palaeoclimatology uses first of all lithofacies methods (Strakhov, 1960; Rukhin, 1962; Nairn, 1964; Schwarzbach, 1950; Sinitsyn, 1967; Ronov and Balukhovsky, 1981; Ushakov and Yasamanov, 1984), whereas prognostic palaeoclimatology undoubtedly prefers isotope and palaeontological methods, which are more indicative of the short-term climatic fluctuations. The estimates of palaeotemperature and particularly of the past precipitation are still rough, including the estimates obtained by the oxygen isotope methods that are more accurate. Palaeoclimatology still needs to be further developed and improved. However, the discussion of these problems is beyond the scope of this book.

Fig. 1. I . Palaeoclimatology and related disciplines.

17

1.2. On the terms “global climatic event”, “climathem”, “climatostratigraphy” Geology uses a “signal”, i.e. cause-and-effect concept of time. The coordination of the events and sediments in space and time is carried out by “markers” that correspond to the greatest extent to the idea of synchroneity. For instance, the causeand-time succession of events such as “volcanic explosion - release of ash - its deposition” represents for a geologist an ideal marker, i.e. a signal about coordination of the events and deposits before and after the volcanic explosion, even though only on a local scale. The drainage episodes of the Bering Strait, which coincided with the emergence of Hipparion and Equus in America, led to a geologically instantaneous (for 100 ka) dissemination of these animals throughout the Old World. Their remains are the best marker for an inter-regional correlation. Fast climatic fluctuations and the dependent changes in sedimentation, organic world as well as geochemical and isotope composition of the remains of animals and plants also represent a rapid cause-and-effect (signal) interaction and consequently a marker for coordinating the geological space and time, including more vague traces of climate change. Climatic signals are valuable because they are global and frequently given. Their shortcoming is that in themselves they are not unique. They can be compared to radiosignals of exact time transmitted every hour, which can be used for checking the right time but are useless for setting a watch on a certain day of the week or a certain year, without additional and relevant information. Therefore, until recently the climatostratigraphic subdivisions such as moraines, loesses, soils as well as spore and pollen zones could have been used only as local units (like the abovementioned ash deposits). Indeed, it is impossible to identify glacial epochs in the Sahara, where ice could not appear at all, or interglacials in Greenland, where ice has constantly been present for millions of years. The attempts to develop any local nomenclature (such as the Alpine, Mediterranean or some other) into a global scheme have always come to nothing. All correlations made proceeding from the apparent similarity in the sequences of moraines, palynological zones and so on turned out after proper verification to be erroneous. Finally, experience and errors allowed one to conclude that such terms as “glacial”, “interglacial”, “interstadial”, “pluvial” etc. themselves are suitable only t o be used locally. Therefore, until now a general climatostratigraphic scale has not been developed either for the USSR or for the entire world. This is the first and the most difficult obstacle in the way of reconstructing the climate of the past (particularly, if the reconstruction is t o be global not local). In order to make a chronological basis for studying climates of the past, it is necessary t o introduce the terms “global climatic event” and “climathem”. It is quite understandable that any local climatic changes in the geological past, be it the advancement of ice, lake-level fluctuations or culmination of pollen of some species in pollen spectra, depend on two factors, namely temperature and atmospheric precipitation, not only annual, but also seasonal. It is just because of this that natural climatic fluctuations turned out to be asynchronous in different places and the time boundaries of the foot and roof of the moraines and pollen zones are transgressive in time. To eliminate this transgression, it seems to be necessary, first,

18

to distinguish in palaeoclimate between changes in temperature and atmospheric precipitation and, second, to use as a climatic reference mark not the entire time interval of changes in temperature and its amplitude but a change in the temperature trend itself. Global climatic events, in the broud sense, are suggested to be understood as the states of the climatic system separated by changes in the temperature trend (or the most pronounced parts of this trend). The empirical evidence shows that changes in the trends in different natural zones on land and in the ocean turn out to be geologically synchronous and are fixed independent of the peculiar features of the local natural environment. We shall come back to the causes of these synchronous events in the Summary. Since climatic events are restored proceeding from the facies appearance of the sediments (actually by their “faces”), the material carrier of the traces of the global climatic event is the stratigraphic subdivision of sedimentary rocks called a “climathem” (Zubakov, 1978), the term being formed by analogy with such terms as the system, erathem, cyclothem. This is a non-taxonomic term, the definition of which implies neither duration nor amplitude of the temperature trend but only synchronous changes. In some particular case temperature changes might even have different signs (for example, against a background of the global temperature increase, certain areas in low latitudes on land and in the ocean might experience some temperature decreases). Thus, in modern palaeoclimatology the term “average climate” should be replaced by the term “global climatic evens”. In this case the local climates can be integrated into the global ones proceeding not so much from their belonging to a common chronostratigraphic subdivision (as in the conventional geological palaeoclimatology), but from the traces of the climatic event, the duration of which might be much less than that of any chronostratigraphic unit. It is natural that here we come across an entirely new methodological situation of correlating the traces of climatic events. That means that the structure of palaeoclimatology will acquire a second box (Fig. I. 1) called climatostratigraphy and climatochronology. Logically it consists of two sub-boxes, one of which solves the problem of correlating the traces of climatic events through stratigraphic methods and the other through chronometric methods. However, in practice both sub-boxes operate in conjunction. Thus, climatostratigraphy is understood as a system of modes, methods and procedures of stratigraphic and chronometric correlation, which uses the traces of the global climatic events as references f o r synchronization.

1.3. On the methods of high-resolution climatostratigraphic correlation and chronological scale of global climatic events The climathems can be used in practice provided two conditions are fulfilled. The first condition requires well-reasoned evidence confirming that the chosen markers at stratigraphic sections (the changes of lithology, facies, fossil remnants) are really related t o a change or rapid growth of the temperature trend. If this condition is

19

fulfilled we can divide with certainty the section into local climatostratigraphic units called kryomers and thermorners (i.e. cold and warm portions of a sedimentary cycle). The second condition requires that distinguished local climatostratigraphic units should have additional information on unique natural features characteristic of a climatic trend, which would contribute to a better identification of the unit in different sections both on regional and global basis. This additional information is called the “characteristics” of a unit in stratigraphy. It allows a climatostratigraphic unit to be transferred from local subdivisions to global units - climathems. These unique features can be of palaeontological nature (the first appearance and the last appearance of guide and representative fossils, the so-called datum levels or datum planes) or they can be geomagnetic (reversions of polarity, or “abnormal” change of declination and inclination recognized by the pattern of curves and loops Table 1.2. Geochronometric dating methods applicable to the Pleistocene ~

Method

Radioisotopic (a) Direct measurement of radioisotopes or decay products Radiocarbon (I4C) Potassium -argon (4nK/ 4nAr) (b) Equilibrium measurement Uranium-series 2 3 0 ~ h 234” /

Organics Volcanics Organic matter, carbonate, coral shell, bone, travertine, tuff

0-350 0 - 200 40- 1000 No limit

231pa/ 235” 234~1238~

4He/U (c) Integrated effects Thermoluminescence (TL) Electron spin resonance (ESR) Fission-track (F.t.) traces of 238u

Chemical (a) Amino-acid racemization Alloisoleucine/isoleucine (D/L) (b) Obsidian hydration (weathering rates) (c) Fluorine - apatite (FCVP) Biological (a) Dendrochronology (growth layers) (b) Lichenometry (growth rates) “Seasonal and event signals” Varve chronology Tephrochronology Paleomagnetic (PM) (a) Incidents of polarity reversals, and geomagnetic excursions (b) Secular variations (c) Viscous magnetization (Irvgrowth)

0-70 N o limit

Quartz, loess, Ceramics, ooze Minerals and volcanic glass

10- 1500 No limit

Organic matter: shell, bone

1 - 1000

Silica-rich lava

I

Bone

1 - 1000

Trees Lichens

0 - 4 (9)

10- 1500

-

100

0-9

Annual deposits Tephra (ash)

No limit No limit

Sediments, volcanics

No limit

Sediments, volcanics Sediments, volcanics

?

0 - 730

20

of changes of the virtual magnetic pole) and radiologic (physical and numerical age of deposits). In other words climatostratigraphic correlation can be successful on the basis of the evidence obtained as a result of section studies by means of different methods. This approach requires large investments, though it pays finally since climatostratigraphy provides the most detailed, highly accurate and truly global (covering land and ocean areas) stratigraphic correlation. The aforesaid demonstrates that the distinguishing of climathems and thus the reconstruction of global climatic events becomes possible only if a large range of different methods of detailed stratigraphy has been introduced into practice, these are micropalaeontological, physical and isotope methods. In fact this is a transition from stratigraphy to high-resolution chronology of climatic events. This allows us to analyse temporal regularities of climatic changes, to investigate their causes comparing empirical chronology of events with orbital rhythms, calculated astronomically. It is apparent that the use of numerical ages would require the development of criteria for validation of estimates. A wide range of chronometric methods is used for the Pleistocene deposits dating at present (Table 1.2). These methods are described in Zubakov (1974), Kaplin (1974), Punning and Raukas (1983), Bowen (1978), Bradley (1985) and Mahaney (1984) many other publications. Each of these methods has its merits and drawbacks. Practice has shown that even the most trustworthy of them, such as the 14C-dating technique, does not guarantee the determination of the true age of the deposits. In fact, the geochemical history of samples dated (buried timber remains, shells, corals) was different in every case. Not a single dated sample represents a closed system, it is subjected to isotope exchange by irradiation of different degrees. Variations of sample numerical ages will often exceed the error of laboratory measurements, resulting in false values. For instance, the most accurate radioisotope technique (14C dating) will yield a series of final datings ranging from 20 to 50 ka for samples, which are obviously older since they belong to the interglacial (the last interglacial occurred 130- 115 ka BP) or to the Pliocene. That is why the numerical ages obtained should be verified against geological, chronometric and palaeomagnetic evidence; this is a necessary condition of the development of a chronological scale of global climatic events (Zubakov, 1974; Symposium . . ., 1982). The geological verification (test) is a very important condition, it is necessary to date continuous sections which have already been comprehensively studied and divided climatostratigraphically with high resolution. It is done to test the validity of numerical age changes along the section, to check for the inversions of age estimates and for the agreement between the obtained ages and data on relative ages of sediments, i.e. the composition of fossil fauna and the like. If a discrepancy is found between the relative age of sediments and the numerical ages obtained, this would give reason t o doubt the validity of the latter. The chronological test is also a very important condition. The sections should be dated by means of various techniques. There is a rule in radiology that a numerical age is valid if close values are obtained by at least two different methods. Some 10- 15 years ago, when 14C dating was used almost exclusively, this condi-

tion could not be fulfilled. Unverified “young” I4C ages made it possible to refer erroneously old sediments like Likhvin sediments in Karukyula and Bolshaya Kosha and very often Mikulino - Karangat - Karginski sediments to the Wurmian interpleniglacial. It has been shown by our experience that the emergence of chronological myths like those of the intra-Wurm interglacial should not be blamed on inadequacy of chronological dating methods but on the use of one particular technique only. A third and very important criterion of the validity of numerical ages is a palaeomagnetic test. Sedimentary rocks are known to maintain the “memory” of such parameters of the geomagnetic field as its orientation (polarity), which changed periodically, and intensity. The technique for measuring the natural residual magnetism (NRM) developed by Khramov (1958), Cox et al., (1964), Harland et al. (1982) and Bradley (1985) allows us to establish not only the NRM polarity of the samples but also many anomalies of the geomagnetic field. There were only a few such anomalies in the Pleistocene (see Chapter 2). Almost all of them were established in their continuing sequence and dated. The transition from the Matuyama reversed polarity epoch ( = magnethem) to the Brunhes normal polarity magnetochron is one of the main signals, which is known to occur 730 ka BP as shown by lava dating by means of the K/Ar method (Mankinen and Dalrymple, 1979). It is extremely important now t o carry out an extensive palaeomagnetic study of all key sections of the Pleistocene used for the development of a climatostratigraphic scale. It is apparent that only these three criteria would make valid the numerical ages obtained.

1.4. The principles of time classification of the global climatic events: Taxonomic differences in the climato - sedimentary cycles and climathems The primary aim of the prognostic palaeoclimatology (and of this work also) is the development of a natural time classification of global climatic events (GCE). And this can be done in only one way, namely through classification of the material traces of the past climatic events, which are found fixed in geological sections. Theoretically these traces could be left by casual catastrophes or occur as more or less regularly recurring phenomena. The former are vividly seen against a homogeneous climatic background, for instance a brief cooling caused by volcanic eruptions or a succession of ice surges. The latter can be seen as a single concurrent succession of events together with other similar preceding and succeding events occurring in a rhythmic and quasi-rhythmic pattern. It is quite clear that these traces are of great value in clirnatostratigraphy, because the rhythmic fluctuations are easily subjected to taxonomic classification. All climatic rhythms and the relevant cyclothems ( = climathems) have a double struc-

’

’

Here, the terms “cycle”, “rhythm”, “periodicity” do not imply a strict recurrence of the same patterns, since in the history of the Earth and climate such strict recurrence is simply impossible. Of course, it is more correct to speak about quasi-cycles, quasi-rhythms and quasi-periodicity. However, to make terminology more simple “quasi-” is henceforth omitted.

22

ture and can be divided into two parts: kryomers and thermomers, or kryochrons and thermochrons. The principles of the structure of sedimentary cycles are a subject of special studies. These cycles can easily be classified by their energy characteristics and get entered into taxonomic schemes according to the time component. In geology, the estimation of the duration in time of sedimentary cycles has been made in different ways, direct and indirect, the former being represented by chronometric methods and the latter by palaeontological and lithological methods. In practice, both approaches are used in combination, which in the majority of the cases allows one t o determine correctly the time succession of sedimentary cyclothems and divide reliably cycles, whose lengths differ from each other by several times. Fast advancement in chronometric methods of numerical dating of the latest sediments and thier magnetostratigraphic correlation (see Chapter 2) has recently given new possibilities for using the data on the duration of past climatic events as a basic for their taxonomic classification. For instance, this has made it possible to compare the length of climatic sedimentary cycles measured empirically with the duration of astronomical rhythms and to look into the time and cause-and-effect relationships between them. Because of these innovations the author (Zubakov, 1978b) was able to suggest using a taxonomic scheme for climathems (Table 1.3) as a basis for time classification of the past climatic events. This scheme has not yet been completed; it includes five units that are most important for palaeoclimatology. Three of them belong to the lower taxonomic rank (nanno-, ortho- and super-clirnathems) and are specially designed for describing the Pleistocene climatic events. They will be substantiated and treated in every detail in the subsequent chapters. The greatest taxonomic unit in the scheme is a trendclimathem. It corresponds t o half the greatest climato - sedimentary cycle known in the Earth’s history, which lasted for about 300 Ma. Meteorology proceeds from a statistical understanding of climate as a combination of different types of weather (Lamb, 1972, 1977; Monin and Shishkov, 1979). The present work is an attempt to suggest an alternative geological definition of climate (palaeoclimate) as a physical essence of a long-term state of the exosphere of the Earth (the atmosphere - hydrosphere and biosphere) resulting from the tectonic evolution of the planet.

1.5. The two climatic regimes in the history of the Earth Some authors when discussing past climates emphasize the principal uniformity of the Earth’s climate since the Late Precambrian, while others claim that it is extremely unstable and variable. This difference would seem to spring from a certain inadequacy of the definition, for in fact if we consider that water and life on the Earth have existed for the last 3.5 Ga and remember that the temperature tolerance of both is very low, then the Earth’s climate appears to be astonishingly stable from the cosmic point of view. However, if the climate of the past is considered through its impact on organism evolution and sedimentation, then only climate variability

Table I .3. Taxonomic rank of climatostratigraphic and climatochronologic units -

~~

.

Content

Climatostratigraphic units ~.

Climatochronologic units, examples

.~

Nannoclimathem (nCT) Orthoclimathem (OCT) Superclimathem (SCT) Hyperclimathem (HCT)

.............. .............. Trendclimathem (TCT)

Parts of 1 .O 2.5 ka climatic rhythm Parts of 90- 100 ka rhythm Parts of 380-320 ka rhythm Parts of 1,200 ka rhythm -

.............. .............. Parts of 300 Ma rhythm

Nannothermochron (Allerad) Nannokryochron (Dryas 111) Orthothermochron (Holocene) Orthokryochron (Warthe) Superthermochron (Waal) Superkryochron (Menap) Hyperthermochron (Tegelen) Hyperkryochron (Pretegelen)

.............. ..............

Trendthermochron (Mesozoic + Lower Cenozoic) Trendkryochron (Late Cenozoic)

w N

24

remains constant, though whether this reflects the variability of local climates or of the global climate is not fully clear. The concept of the alternation of cosmic winters (according to Chumakov (1984), glacial eras) with cosmic summers (thermal eras) is generally accepted. This is an empirical fact, one of the greatest achievements of paleocIimatology of the 1940s. While the temporal distribution of glacial geological indicators such as salt accumulation, coal formation and other paleoclimatic events can be equally attributed to regional paleogeographic changes and to the changes of global climate (Fig. 1.2), MEAN GLOBAL TEMPERATURE present

cold

my

warm

MEAN GLOBAL TEMPERATURE

present wet

dry

I Quo t e r na r y Pliocene

Oligocene

F Paleocene

65

1Jurassic

W.9

k-L

Cambrian

1000 -

2000 -

3000-

4000-

Fig. 1.2. Generalized temperature and precipitation history of the Earth. The curves are drawn to represent postulated departures from present global means, but only relative values are indicated (after Frakes, 1979, fig. 9.1).

25

the obvious question is whether the mean global air temperature has changed and, if so, what the change amounts to since the mean global air temperature is a major parameter characterizing the physical state of the global climatic system as a whole.2 The question appears to have two answers, in two groups of hypotheses. The first answer is given by Lisitsyn (1980): “Climatic belts have always existed . . . The mean Northern Hemisphere temperature remained virtually unchanged (ranging from 17 to 20°C) in the past 200 Ma, meaning that no global warming or cooling occurred in the Mesozoic and Cenozoic (including during the Quaternary Ice Age). The distribution of surface heat and moisture has changed though, caused by changes of the heat exchange mechanism (Monin, Shishkov, 1979) . . . Thus, major climatic changes of the geological past are believed to be determined by the development of the ocean and sub-polar land-masses in the Northern (and Southern) Hemisphere. When the Earth’s poles were on land, polar continental ice sheets would necessarily develop (Monin, 1977, p. 198). The times when land masses clustered near the poles corresponded to glaciation periods . . .” (Lisitsyn, 1980, pp. 391 - 393). A large number of geologists accept this hypothesis, particularly about pre-Cenozoic glaciations. A uniformitarian and Earthy concept of causes of climatic change is part and parcel of these paleogeographical theories. The second answer involving significant repeated changes of the mean global air temperature (nearly doubled according to theory) is given by proponents of two other groups of hypotheses: (1) one of them allows for temporal variations of solar radiation caused by external factors such as changes of the solar constant, passage of galactic dust clouds, collision with comets, supernova star flares and so on (Balukhovsky, 1966; Salop, 1977); ( 2 ) the other assumes changes of the Earth’s “heat-keeping” mechanisms (i.e. greenhouse effect, albedo and the like) as determined by internal factors, primarily changes in atmospheric gas composition (Arrhenius, 1903; Chamberlin, 1899; Budyko, 1972, 1980, 1984; Manabe and Wetherald, 1980; Berger and Crowell, 1982; Barron, 1983, 1984, 1985 etc.).

The present authors belong to this latter group of researchers; they will attempt in this chapter to summarize current evidence of pre-Pliocene climate changes and to show that such evidence indicates significant variations of the mean global temperature the extreme values of which can be interpreted as principal qualitative changes in the working regime of the climatic system itself. This approach has only become possible recently on the basis of the quantitative paleoclimatic record of isotopic and bioclimatic curves obtained by deep-sea drilling. The study of interaction mechanism in the atmosphere - ocean - kryosphere system has also contributed to this approach. We mean here studies made within the framework of DSDP, CLIMAP, CENOP and other projects. Particularly im-

’

The term “mean global air temperature” in physical climatology means the temperature of the Earth’s surface averaged over latitudinal zones, with consideration of their areas.

26

portant are recent publications on the role played by thermohaline water circulation and deep sea currents in the formation of the global climate (Weyl, 1968; Kroopnick et al., 1977; Schopf, 1980; Seibold and Berger, 1982; Brass et al., 1982; Fischer, 1982; Kennett, 1977; Hay, 1983; Barron, 1983, 1985; Ciesielski et al., 1982; Keller, 1983a,b and many others). The amount of incoming solar radiation is known to depend on latitude, while albedo is dependent on the properties of the underlying surface. The ocean (from 30"N to 30"s) is the main accumulator of the solar heat since the sea surface has heated up t o 33°C which is the equilibrium temperature, determined by the specific heat capacity of sea water and a negative feedback m e ~ h a n i s mThere .~ is a constant heat deficit in high latitudes, so the climatic engine has to redistribute heat over the latitudes and between the oceans and continents. Excessive heat from the tropics is transferred towards the poles as warm air masses, warm sea currents and as the latent heat of water vapour condensation. This mechanism is faily well studied both by climatologists and oceanologists. It is assumed that this mechanism has been always acting though with three major variables: (a) changes of astronomical parameters, determining the short-period rhythms of climatic processes; (b) changes of the paleogeographic environment meaning the latitudinal distribution of land masses and oceans, changes of their extent, changes of the topography of the Earth's surface and oceanic currents; (c) dynamics of the atmospheric and surface ocean circulations which is believed to be of an arbitrary nature mainly determined by winds. A fourth factor was discovered only recently: it is the thermal state of the deep ocean, and the water-mass mixing which appear to regulate the long-term state of climatic system. Heavy bottom waters are known t o be formed by two mechanisms - a strong salinity increase typical of water basins with evaporative processes a condition now observed in the Mediterranean Sea and Persian Gulf (Brass et al., 1982; Hay, 1983, and others); and with a cooling of normally saline waters in sea-ice regions, such as occurs now in waters adjacent to Antarctica (Weyl, 1968; Ciesielski et al., 1982; Johnson, 1983 and others). Since the first mechanism is typical mostly of marginal low-latitude seas, sinking water would have very high salinity and high temperature, while the second mechanism brings down cold water. The presence of a freshened surface layer results in a strong density stratification, which prevents vertical water circulation. That is why no bottom water is formed in the Arctic Basin. At the present time dense deep waters are formed in the Mediterranean Sea which is the source of Warm Saline Bottom Waters (WSBW), the volume of which is estimated to be equal to 1 S V .North-Atlantic ~ Bottom Waters (NABW), about 1 Sv in volume, are formed in the Greenland and Norwegian Seas. Dense deep waters form also in Antarctic waters. It is here that the major bottom water volume is formed (7 Sv), including very cold and dense waters near Antarctic coasts (AABW) and the Antarctic

' Negative feedback manifests itself in typhoons which mix surface and deep waters of the ocean; such typhoons occur only in areas where the sea surface temperature exceeds 26°C (Schopf, 1980). 1 Sv (Sverdrup) corresponds to the present day run-off of all the world's rivers, being lo6 m3/s.

27

Intermediate waters (AAIW) which are less dense in the Antarctic Convergence Zone. Comparing the above with the paleogeological evidence obtained by the analysis of deep-sea cores which yield a systematic and continuous record of paleotemperature (covering almost 120 Ma), oxygen-isotopic data (Douglas and Savin, 1973; Savin et al., 1975, 1977, 1981; Shackleton and Boersma, 1981; Blanc et al., 1983; Keigwin, 1979; Keigwin and Keller, 1984; Leonard et al., 1983; Loutit, 1981; Shackleton and Cita, 1979; Duplessy, 1981; Thunell, 1979, and others) and data inferred from the paleontological record by factor analysis (Barron and Keller, 1982, Barash et al., 1983; Berggren, 1978; Ciesielski et al., 1982; Ciesielski and Weaver, 1974; Diester-Haas, 1979; Kennett, 1977; Sancetta et al., 1983; Cita et al., 1977; Keany, 1978; Poor, 1980; Prell et al., 1980; Ruddiman, 1971, Keller, 1981, 1983a,b, and others). We may conclude that: (1) deep-sea sediment sections store invaluable information about the mechanisms and history of the world's climate, as a system of long-term interactions between the atmosphere, ocean, kryosphere lithosphere and biosphere; (2) the temperature of the deep ocean can be used as a quantitative parameter, reflecting changes of the global thermal regime during prolonged intervals of time; (3) long-term trends in the changes of the ocean's thermal state correlates with climate indicators, these were previously used by geologists to develop a theory about alternation of glacial eras and thermal eras over the course of geological history. It appears that the beginning of the last glacial era is concurrent with the formation of the so-called psychrosphere in the ocean (Charnberlin, 1899; Keller, 1983) which is cold heavy bottom waters with temperature less than 8 - 5"C, while the last thermal era corresponds to a time of warm botton-water accumulation and hence of an oceanic circulation radically different from the present one; (4) In the process of ocean water stratification and ocean sedimentation certain stages or steps develop, divided by various dramatic events such as catastrophic increases in bottom current speeds, sea level changes, isotopic shifts and changes in the depth of carbonate compensation (CCD), changes of sedimentation rates and sediment composition, changes of fossil microfauna, and also variations in the CO,/O, ratio. Accurate dating of these stages by means of paleomagnetic, chronometric and biostratigraphic records gives a new insight into the history of the world's climate and help to better understand causes and mechanisms of its changes.

Paleo-oceanographic evidence on the climates of the past is in fairly good agreement with the Arrhenius - Budyko theory that CO, determines the climate evolution. According to this theory a high C 0 7 concent in the atmosphere would determine a warm tlimate, whereas a low CO, content would indicate a cold climate (Arrhenius, 1903; Chambedin, 1898; Budyko, 1980, 1984; Manabe and Wetherald, 1980; Manabe and Bryan, 1985; Manabe and Braccoli, 1985). On the basis of calculations of igenous rock volume by Ronov (Budyko et al., 1985), Budyko (1986) estimates that the global air temperature increased during 26 major Phanerozoic stages, as compared to the modern mean global temperature (14.2- 15OC). He assumes that a doubling of the current CO, level (0.034%) would cause a mean

28

global surface air temperature increase of 2.5”C also taking into account changes of the solar constant with time and albedo differences due to variations in the ocean surface and the formation of ice sheets. Budyko’s calculations lead to the conclusion that there have been repeated occurrences of the “greenhouse effect” or “greenhouse climate” in the past. In the Devonian, Lower Permian and Mesozoic mean air temperature exceeded the present one by 5.4 - 10.6”C, meaning it was about 20 - 25°C. These estimates agree with those other investigators have obtained by different methods; for instance, by paleoclimatic mapping, Sinitsyn (1965) found the mean air temperature in the Cretaceous t o be 26”C, while in the Albian - Cenomanian (1 13 - 92 Ma) according to Barron (1983) it ranged from 20.5”C to 28.5”C. Isotopic analysis of the belemnite rostra has yielded water temperatures in the Albian as 21”, in the Cenomanian as 16.5”, in the Coniacian and Santonian 20” and in the Maestrichtian 16°C. On the other hand, sea surface temperatures in lower latitudes decreased by 2 - 4°C during the height of the Wurm glaciation and annual air temperature in the middle latitudes dropped by 5 - 8°C (CLIMAP Project Members, 1976). Thus, the amplitude of mean global temperature variations from thermal eras to glacial eras reaches 10- 20°C. The present authors believe that changes of global climate with such an amplitude and a duration of hundreds of millions of years differ in principle from climate variations of shorter duration and smaller amplitude, because then the entire mode of operation of the climatic engine changes, which in turn would generate major changes in organic life, as life adapts to a new mode of climatic engine operation. That is why the authors in accord with Fisher (1982) think it worthwhile to introduce the concept of the climatic regime as a qualitative state of the Earth’s climatic system. During the past 2.7 Ga this regime has changed five times. A glacial regime existed 2.6 - 2.2 Ga, 770 - (940) - 620 Ma, 450 - 400, 330 - 240, and 38 - 0 Ma, each time continuing for not fewer than 50 Ma. A description of these events is given in an interesting book Winters of our Planet by John et aI. (1979). The greenhouse effect has also ocurred not fewer than five times, continuing for not less than 150-200 Ma. Fig. 1.3 shows how the greenhouse effect coincides with an intensification of volcanic activity and marine transgression, while glacial regimes coincide with marine regression. This allows the alternations of regime to be associated with cycles of mantle processes which caused by formation and destruction of Pangaea. A rapid increase of atmospheric CO, levels makes the problem of the alternation of glacial and greenhouse regimes in the Earth’s history worth more detailed consideration. 1.5.1. Main features of the glacial climatic regime As seen in Fig. 1.4 the most pronounced thermal stratification of the ocean waters may be dated as 38 k 1 Ma, when bottom water temperatures decreased to 6 - 4°C instead of 17 - 13°C of the Paleocene. This temperature is agreed to be taken as a boundary between two climatic regimes, although typical features of the greenhouse

29

regime start to disappear earlier, approximately 50 Ma, when the NABW formation began (Berggren and Hollister, 1974; Blanc et al., 1980). While the typical features of a glacial regime formed much later, about 14- 10 or even 7 . 7 Ma for which reason this broad transient interval of 50 to 7 Ma will be discussed later. I f the modern climate which corresponds to an average state between the climate of the glacial epochs of the Pleistocene and Pliocene and the thermal optima of the Miocene and Pliocene is considered to be a model of the glacial regime, then its most typical features would be:

(I) The poles of the Earth should lie on one of the continents with most of the land masses located in high and middle latitudes so determining a climate asymLong-Term Climatic Oscillations Recorded in Stratigraphy

1

6 x

lo8

Y.5 P

4

2

0

Fig. 1 . 3 . Relation of inferred climates to secular in volcanism, sea level, and organic diversity. Volcanism: emplacement of plutons in North America, after Engel and Engel (1964); sea level: A, firstorder eustatic curve of Vail et al. (1977), B. Compromise between North American and Russian records, constructed from Hallam (1977); the scale as left refers to this curve. Biotic record: N, Stehli et al.’s (1969) curve of disappearance of animal families, C, net gain- and loss curve of Cutbill and Funnel (1967), overlap shaded. Inferred climatic states from Fischer (1981); minor oscillations (which may bring about growth of ice sheets, shaded) after Fischer and Arthur (1977). From Fischer (1982, fig. 9.3).

30

metry of the hemispheres and a large contribution by albedo to the climate variations of the continental hemisphere. (2) The existence of a kryosphere including surface, underground and marine glaciation. (3) The existence of a psychrosphere closely related to the kryosphere. (4) Temperature asymmetry of the ocean and atmosphere, mean ocean surface temperature being 5.7"C, which is more than twice as low as that of the air, which is 14.2- 15°C. (5) Ocean waters rich in C 0 2 and a C02-depleted atmosphere because of the higher solubility of C 0 2 in cold water. ( 6 ) A temperature gradient between the equator and high latitudes in the ocean and between the surface and bottom which in the tropics reaches 15 - 20°C. (7) High-speed vertical ocean circulation, a short time of water exchange (from 250 to 100 years according t o Hay, 1984) high oxygen content of the waters, highspeed bottom currents (up t o 0.5 m/s) flowing from high latitudes to the equator, and widespread upwelling. (8) An atmospheric temperature gradient between the equator and poles; now it

Fig. 1.4. Oxygen isotopic paleotemperature data obtained from analyses of planktic and benthic foraminifera (and some nannofossils) from DSDP cores. 1 - Site 47, Northwest Pacific; 2 - Site 55, Western Equatorial Pacific; 3 - Site 167, Equatorial Pacific; 4 - Site 357, Centrai South Atlantic; 5 - Sites 277, 279 and 281, mid-latitides of the Southern Hemisphere, Campbell Plateau; 6 - Sites 277, 279 and 281, high latitudes of the Southern Hemisphere (benthic data). Data for the last 60 Ma were compiled by Borzenkova (1981a) from Douglas and Savin (1971, 1973); Shackleton and Kennett (1973); Savin et al. (1975); Boersma and Shackleton (1977); Savin (1982). Data for the time interval from 120 to 60 Ma were compiled by Krasheninnikov and Basov (1985) from nanDouglas and Savin (1973), Savin et al. (1975); 7 - Central Pacific; 8- 10 - North Pacific, 9 nofossils; 10 - benthic foraminifera. 1 - greenhouse - thermohaline regime; I1 glacial psychrospheric regime; I11 transitional state. ~

~

~

~

*

31

is 28.2"Cin summer in the Northern Hemisphere while in the Southern Hemisphere it is 40.2"C in summer and 74.4"C in winter. (9) Wind-controlled atmospheric circulation with strongly developed cyclonic processes, meridional transport and prevailing westerlies. (10) Strong wind-induced oceanic currents, including the Antarctic Circumpolar Current, which thermally isolates Antarctica. (1 1) Well-pronounced zonality over the land and over the ocean. (12) High sensitivity of the high-latitudinal climate to changes in solar radiation controlled by the Milankovich mechanism, and distinct rhythms of climatic variations with cycles of 41, 100, 400, 1200 ka.

Fig. 1.5. Geographical - climatic zones throughout the glacial - psychrospheric regime - Late Cenozoic glacial maximum (from Chumakov, 1983). Glacigenic sediments: 1 - till; 2 icelaid drift; 3 - marine glacial drift; 4 - coal; 5 iron and evaporites; 7 - marl; 8 bauxite and laterite. manganese ores; 6 of warm Vegetation: 9 - of tundra; 10 - of cold rteppe; 1 1 - of cold temperate zone; 12 temperate zone; 13 - thermophilic. Zoo indicators: 14 - warm-loving tetrapods; 15 coral reefs. Climatic belts: 16 - glacial and periglacial; 17 - temperate; 18 - warm non-tropical; 19 - arid and Temi-arid; 20 - tropical rain forest. ~

~

~

~

~

~

32

Figure 1.5 shows the paleogeographic situations in the last glacial regimes. There is abundant evidence of the similarity between the rhythms of Pleistocene climatic events and other glacial eras. For example, Andersen (1982) made a statistical study of the stratification of the evaporite sequence of the Castile and Bell Campon formations (Texas) which revealed distinct climatic rhythms with duration of 0.4 - 0.6; 1.1 - 1.4; 2.2 - 2.7; 17 - 25 and 100 ka. Spencer recognized 47 glacial horizons combined into 17 glacial periods in a 870-meter section in his detailed study of Late Proterozoic tillites (Scotland) (John et al., 1979). There is an abundance of similar data (Zubakov and Krasnov, 1959; Strakhov 1962; Frakes, 1979; Monin and Shishkov, 1979, and many others). 1.5.2. Main features of the greenhouse clinzatic regime

The last ice-free period or Siberian Thermal Era, as it was called by Chumakov (1984), continued for about 200 Ma encompassing the whole of the Mesozoic and half of the Paleogene. Obviously the Earth’s climate experienced different changes, alternating from humid to arid and from equable to hot (Fig. 1.2). However, the performance of the climatic engine remained unchanged. Geologists consider the Mesozoic climate to be “normal” contrary to the “abnormal” present climate (Schwarzbach, 1950; Strakhov, 1962, and others). Let us examine the main features of such a climate regime using as an example a well-documented 10 Ma period from the Late Paleocene (Thanetian) to the Early Eocene (Ipresian), i.e. 60.5 - 50.5 Ma BP. We shall draw extensively on the records of Sinitsyn (1965, 1967, 1980) for the USSR, of Golbert et al. (1977) for Siberia, Wolfe (1980) for North America and Frakes (1979) for the Southern Hemisphere; we shall also use new data on Ellesmere Island (McKenna, 1980), Hickey et al., 1983), Baranova and Biske (1979) and Biske (1981) for the north-eastern USSR. It should first be noted that the paleomagnetic data (Barron, 1983, 1985; Hickey et al., 1983) show the distribution of land masses to be roughly the same as at present, that is paleolatitudes d o not differ from the present ones by more than 2- 5 ” . The Iand/sea distribution (Barron, 1983, 1985) and land surface topography are also shown t o be virtually the same. Whereas sedimentalogical evidence shows the Early Paleogene climate to have nothing in common with the present climate. The Early Paleogene sediments, with typical thick (up to 40 meters) caoline weathering crusts and highly carbonaceous formations were found to occur in high latitudes, extending to the Polar Sea coast (70”N). Bauxites formed in Siberia, low-latitude laterites and ferrallitic soils are found to have very distinct sections and to be more widespread than at present (Sinitsyn, 1965, 1980; Strakhov, 1962; Golbert et al., 1977). Marine sediments also display numerous warm climate indicators; these are sandy-glauconites and abundant fossils rich in large tropical molluscs, nummulites and reef-forming corals. Ivanovsky (Sinitsyn, 1980) reports on solitary corals Leptoria and Sphenophyllum found as far north as the West-Siberian Sea (60”N). Opal lime mud deposited within the Polar Basin, which at present is typical of upwelling zones with high productivity (Clark, 1982).

33

Fig. 1.6. Geographical climatic zones throughout the greenhouse- thermohaline regime (after Chumakov, 1983). For explanation see Fig. 1.5.

-

Early Eocene

Terrestrial fossil flora are even more informative (Fig. 1.6). Tropical rain forests were widespread on the Atlantic Coast flatlands with highly diverse taxa (flora of London clays for 70 - 80% are made up of plants that became extinct in the Eocene; they are known to belong forma-genera of Stemma Normupolles, Postnormupolles and others. Genera which were not extinct included numerous palms and Ficus, sandalwood, Quaiucum and various tree-ferns and lianas. Tropical mangroves with Nipa palms ranged into coastal areas. Forests equivalent to the latter are found now in Vietnam and Hainan where they grow in environments with I,- . . . 16"C, I, . . . 26"C, Ian . . . 20"C, annual precipitation sum of 1000- 1500 mm (Sinitsyn, 1980). Rain forests were changed in the east by seasonally wet forests with Normupolles stemma dominating among Myrtus, Myrica, laurels, magnolia, platanos evergreen oaks and the like. The region of the present-day asiatic deserts and semi-deserts was covered by savanna and steppe with flatland forests including forma-genera, and Myricu, Tuxodiuceue, tree-ferns, gum-trees, palms and the like. Coniferous forests with cedars and Aruucuria inhabited flat interfluves. Taxa requiring less than 500 mm of precipitation as an annual sum did not grow in Asia at that time (Sinitsyn, 1965). Golbert et al. (1977), on the basis of palynological evidence, reconstructed hot, humid or seasonally wet climates in Northern Kazakhstan with Ian . . . 22-25"C, annual amplitude of 4-6°C. Such a climate would be

34

characterized by annual precipitation of 800 - 1200 mm, the dry season continuing for 3 - 5 months. Forests in high latitudes appear to be even more astounding. As shown by Sinitsyn (1965,1980)and by Golbert et al. (1977),Wolfe (1980),Baranova and Biske (1979), Biske (1981)there was a circumpolar zone of conifers and broad leaved forests in the Northern Hemisphere, half of this vegetation consisting of extinct formal genera. This zone is believed to extend northward to the Polar Ocean, that is 70-76"N. The vegetation for one third was found to consist of Myricu, Myrtus, gum-trees, Engelhardria, laurels, tsuga, palmae sabal, holly groves and Ficus, that is species whose descendants are not found outside the sub-tropics at present. The other two-thirds extended as far north as New Siberian Islands, Spitsbergen, Ellesmere Land and the Axel - Heiberg Islands. Species that inhabited those forests growing in close proximity are now found in three different geographical zones, that is in tropical, sub-tropical and temperate zones. That was a complex polydominant flora which has no modern equivalents. It is similar though to that of southern Japan and Middle China, where the winter temperature is 8- lO"C, in summer it is 22- 25"C,the mean annual temperature is 18" with the annual precipitation sum being 1,000-2,000 mm (Sinitsyn, 1980). As shown by Kemp (1978)similar fossil floras are found in southern high-latitudes (60- 70°S), in Australia, where palms were also found (Frakes, 1979). All authors agree that the high-latitude climate was equable and humid, with mild essentially frostless, winters, extremely favourable for coal accumulation. The organic life characteristics in the Thanetia- Ipresian would not be complete without paleofaunal data. A large widespread group of Dinoceras with numerous paleoungulate animals (Prodinoceras) and tapiridians (Homogalax, Hyracotherium) inhabited savannas. Hippopotami (Coryphodon), typical inhabitants of swampy forests, were found in the section in the deserts of Gobi and Central Asia (Sinitsyn, 1980; Demberlane 1984). Estes and Hutchison (1980)and McKenna (1980)reported on very interesting findings o f the fossil vertebrates in the sections of Eureka Sound formation (Ellesmere Island, 78"N). These sections display an ample amount of fossil rodents, salamanders (Piceoerpeton), monitor lizards, snakes, turtles, tortoises (Geochelone), crocodiles (Allognatsuches), lemurs and other animals as well as birds and fish, equivalent t o those inhabiting present-day tropical and sub-tropical forests, since they do not tolerate temperatures below 10 - 12°C.Isotopic foraminifera1 analyses from marine units of these sections suggest a water temperature as high as 15°C (McKenna, 1980). Findings of the fossils of evergreens along the coast of the Polar Ocean were always controversial ever since Svalbard was explored. Discoveries of fossil tropical fauna (alligators, lemurs, monitor lizards) at 78"N have added to this effect. Strakhov (1962)and Wolfe (1980)argue that evergreens could dwell in the polar night environment. They attribute their occurrence in the fossil fauna and flora of high latitudes to a lower tilt of the Earth's spin axis to the ecliptic (5 - 15°C).Barron (1984)modelled orbital variations with the spoon axis inclination of 5" and 15",his simulations showed no polar night in high latitudes, and a drastic decline in the solar radiation income and hence a temperature decrease, while in the lower latitudes an

35

increase of the incoming solar radiation would bring about a temperature rise up to lethal values (35 - 34°C). Sinitsyn (1980) suggested an increase of the solar constant as a possibility which is equally inacceptable. Circumpolar development of paratropical plants (found even at New Siberian Islands, 75 - 80”N) refutes Donn’s (1982) conclusion on the erroneous paleolatitudes of Ellesmere Island. Thus, we cannot help admitting a paradox of tropical flora and fauna in the lightlimited environment (3 - 5 months of polar night). We have to agree with Krasilov (1985) and Sinitsyn (1980) who suggested the possibility of adaptation of evergreens to seasonal illumination. The lower rates of metabolism of hardwood plants during the polar nights would explain the annual ring formation in the warm humid climate (!).

Hickey et al. (1983) have given new insight into the problem. They carried out detailed magnetostratigraphic examination of the Eureka Sound Formation section which is more than 3 km in length and compared it with the Lowrie - Alvarez (1981) scale of magnetic anomalies. Thus, all units of the section were accurately dated. This allowed them t o make chronological comparison of the first appearance of floral and faunal taxa in the Eureka Sound formation section and in the lower latitude sections. It appeared that many vertebrate taxa (Hyrachyus, Pantolestes and others) inhabited Ellesmere Island some 2 - 4 Ma earlier than the lower latitudinal areas of North America. Fossil floras display even more astounding irregularity in time. Spores and pollen of numerous taxa (e.g. Pisrillipolenires mc gregorii, Mefasequoia occidentalis, “Carya” antiquorum etc.) found in the Eureka Sound formation section are 18 Ma younger than similar taxa found in southern North America. It is thus established that many faunal and floral taxa (Terraclaenodon, Hyracotherium, Plagiomene, etc.) are of polar origin (Hickey et al., 1983). This confirms the inference (Krasilov, 1976) that the polar latitudinal areas with their singular combination of warm climate, seasonal illumination and high productivity were centers of evolution of flowering plants and vertebrate taxa. Thus, we can firmly conclude that the only possible explanation of these features of organic life and sedimentation in the Early Paleogene both in high and low latitudes is a principal difference of the Early Paleogene climate from the present climate. The Adem model, basd on a uniformitarian concept of climatic engine performance, gives no explanation, since it implies that winter temperatures in high latitudes would have been below 0°C (Donn, 1982) which is contrary to the available evidence. This can only mean that the uniformitarian approach cannot be used in the description of the Early Paleogene climate. Obviously the Paleogene climatic engine performed differently, its regime has been defined, following Budyko (1980) and Fisher (1982), as the greenhouse regime. The comparison of isotopic deep sea drilling data published by Douglas and Savin (1971, 1973), Savin, 1982; Savin et al. (1975), Kennett (1977), Boersma and Shackleton (1977) and others (Fig. 1.4), with the aforesaid allows us to reveal major features of the greenhouse climate: (1) Neither kryosphere, nor psychrosphere existed. (2) Formation of the Warm Saline Bottom Waters (WSBW) took a longer time

36

geologically due to the widespread development of epicontinental marginal seas in lowerer latitudes resulting in a slight increase of the mean sea temperatures accompanied by a decrease of the temperature gradient to 5 - 10°C from the equator to the pole in the tropics (from the water surface to the bottom). (3) A decrease of the CO, dissolved in the ocean waters and an increase of the CO, levels in the atmosphere by 5 - 10 times as compared to the present levels would lead to enhanced greenhouse effects and a rise of the mean global temperature to 25 - 27°C. (4) A decrease of air temperature gradient between the equator and poles. In winter it is about 20”C, resulting in a weakening of the atmospheric centers of action and a drastic decrease of the effects of the westerlies and the equator-to-pole air mass transport. ( 5 ) The development of a certain deep-water mechanism redistributing heat between latitudinal zones. The heat mainly accumulated in middle latitudes is transferred polewards by deep ocean currents, its amount being much larger than that transferred by winds in the present day despite low velocities of ocean currents. In high latitudes warm deep waters would have experienced upwellings as a result of their salinity decrease. Since polar land masses would cool faster than the Arctic Ocean during the polar night, winter offshore winds would cause warm water upwellings along the Arctic Ocean coasts. (6) Thus, the climatic engine in the greenhouse environment would accumulate heat in the deep ocean and keep the atmosphere warm, “pumping” it from middle to polar latitudes by means of deep currents. (7) This mode of operation of the climatic engine would inevitably lead to isothermal climate with small amplitude of seasonal air temperature variations at all latitudes, with enhanced cloudiness and accelerated turnover of atmospheric precipitation. (8) Climatic zoning is not well pronounced. There would be only three zones: (i) an equatorial year-round humid zone with ferrolite weathering and temperatures 2-3°C higher than the present ones; (ii) tropical seasonally wet savanna with lateritic weathering and annual precipitation not less than 500 mm; (iii) para- or quasitropical high-latitudinal zone with kaoline weathering and intense coal accumulation. Within this zone an area with polar monsoon climate having no present-day equivalents should be distinguished (polar quasitropics). It would have seasonal illumination and frostless and snowless winter. We can further elaborate upon this. First, the polar quasi-tropics with a distinct annual rhythm should be highly sensitive both to changes of paleogeographic environments and to greenhouse effects, thus the area would be most favourable to the evolution of organic life. Second, the precession rhythm of 19 - 23 ka seems to be the most important factor among other Milankovitch orbital mechanisms affecting climate under greenhouse regime. Third, small amplitudes of seasonal temperature variations would open t o injury the organic life typical of greenhouse climate by external impacts whether these are “impact catastrophes” (Budyko, 1984; Budyko et al., 1986) or “injection events” (Thierstein and Berger 1978). Fourth, the alternation of climatic regimes in the Earth’s history can be explained

37

only by the records of historic geotectonics, that is periodic changes of land/sea distribution and finally a transition from one climatic regime to another is believed to be a prolonged and complex process taking tens of millions of years. Thus a comparison of long-term climate-forming features in the time intervals of 40 - 30 and 60 - 50 Ma BP has demonstrated that these features would form only under the assumption of opposite thermal trends wich continued for tens and hundreds of million years. These trends, however, are more typical of the ocean than of the atmosphere. They are manifest mainly in the accumulation of highly saline bottom waters in the Mesozoic-Eocene and of cold heavy waters in the Late Cenozoic. It is these processes rather than volcanic activity that would determine high CO, levels in the atmosphere in the Mesozoic - Eocene time and CO, continued to decrease in Late Cenozoic (Fig. 1.7). Fisher (1982) defined the foregoing climatic environments as the climates of glacial and greenhouse types attributing them to volcanic activity. This terminology seems inadequate. First, these environments are not the climate types, they represent taxonomically something bigger. These are two different modes of operation of the Earth's climatic engine. Their appearance is associated with profound tectonic reconstructions in the mantle and Earth's crust which might be triggered by changes of the galactic orbit. Second, the glaciation and greenhouse effects seem to be not causes but consequences of the two operative modes of the climatic engine. Since the ocean is the main element of climatic engine, it seems more appropriate to call the described regimes in climatic oceanic terms, that is a psychrospheric regime and thermohaline regime. The name, though, is not that important; we should bear in mind that these regimes correspond to the largest stages of the global climate evolution in the geological history of the Earth. They represent the largest taxonomic unit 79:

r

1

I

1 80%

8

I

60

I

1

40

1

1

20

1

1

0

I

I

20

I

I

W

I

I

60

I

I

I

80"s

Fig. 1.7. Compared mean annual air temperature during greenhouse - thermohaline and glacial psychrospheric ocean - climate regimes. 1 - Late Cretaceous (after Barron, 1983); 2 - Pliocene optimum (after Borzenkova and Zubakov, 1985); 3 - present time (after Rubinshtein, 1970); 4 - Late Wurm (after CLIMAP members, 1976; and Borzenkova). ~

38

of paleoclimates and the authors suggest t o call them trendclimathem, thus introducing a new term (Fig. 1.7).

RCsumC (1) The existing pattern of “mean climate” (paradigm) puts certain constraints on paleoclimatic reconstructions due to time limits of the chronostratigraphic scale neither allowing a detailed study of history of climates to be made, not to identify cause - effect relationships of climatic changes. (2) The suggested alternative pattern (paradigm) of global climatic events (GCE) seems to be more adequate for the identification of climatic signals of high accuracy, these signals are represented by rapid changes of temperature trends simultaneously over the whole globe. This new concept is based on climatostratigraphic evidence and methods of comprehensive climatostratigraphic and climatochronological high-resolution correlation. (3) A meaningful periodical classification and terminology of paleoclimates and global climatic events are needed to describe and study the history of climate. These can be developed on the basis of empirical climatostratigraphic classification of sedimentary cycles, i.e. climathems. (4) The authors suggest that the main unit of geological classification of the paleoclimates, a trendclimathem, should represent alternating climate regimes that is greenhouse - thermohaline and glacial - psychrospheric regimes which together form a climatogeological cycle covering about 250 - 300 Ma.

Chapter 2

DEEP-SEA STANDARD FOR GLOBAL CLIMATIC EVENTS 2.1. History of climatostratigraphic study of the Pleistocene

The term Pleistocene was introduced by Lyell (1830- 1833, 1840) for the last epoch of the Tertiary, for the Age of Man, and for the deposits which contain about 90% of molluscan species still living. The Pleistocene roughly corresponds to the Diluvium established by Buckland (1823). As early as 1846 Forbes showed that in fact the recognition and division of the Pleistocene can be based only on climatostratigraphic data. He also proposed a new interpretation for the term Quaternary Period’ as the time encompassing the Glacial - Diluvial- Pleistocene Epoch and post-Glacial -Alluvial, or Holocene Epoch in the sense of Forbes. This unhappy confusion of terms and meanings for the purpose of classification led to still lasting misunderstanding of the extent and position of the lower boundary of the Pleistocene and Quaternary. Here we accept a traditional position of the lower boundary of the Pleistocene proposed by Gignoux (1910) and Pavlov (1925) and placed at the base of the Sicilian regional stage of the Mediterranean (1.15 Ma according to Ruggieri et al., 1983) within the marine sections. On land, the boundary corresponds to the top of the Villafranchian and is marked by the appearance of mammalian fauna of the current generic composition (1.0 f 0.1 Ma according to Azzarolli, 1983), the appearance of the genus Homo (1.2 Ma), and the base of glacial tills of the first cycle of continental glaciation in Europe (below the Jaramillo event after Zagwijn and Doppert, 1978). In the USSR this boundary coincides with the base of the Chaude beds in the Black Sea basin (1.1 Ma according t o Zubakov et al., 1975, 1977), and the base of the Tyurkyany Formation of the Caspian Sea basin (0.95 - 1.07 Ma according to Ganzei, 1984). The close agreement between the above dates gives evidence for the stratigraphic validity of the traditional Pliocene - Pleistocene boundary, whose position was unambiguously determined in different parts of the world long before numerical estimates were used.2 Effects of climatic changes have been used as a basis for detailed stratigraphic subdivision and correlations of Pleistocene ( = Quaternary) deposits from the very beginning of their study. However, the effects were recognized through studies of lithogenesis, geomorphology, migrational changes in fossil assemblages associated with certain environments, etc. As a result, the climatostratigraphic nature of local

’

This was initially proposed by Desnoyers (1829) to embody the current Middle and Late Miocene, Pliocene, and Pleistocene, i.e. the last 15 14 Ma (Leonov, 1973 - 1974). Recently some workers proposed to lower the Pliocene - Pleistocene boundary and place ii below deposits 1.6- 1.8 Ma old (Nikiforova et al., 1982; Ruggieri et al., 1982) as recommended by the INQUA Subcommission (Aguirre and Pasini, 1985) or at the level of the deposits with an age of 2.5 - 3.0 M a (Zagwijn, 1974; among others).

’

~

40

and correlative Pleistocene units has remained uncertain. They were mainly recognized as regional stages, horizons, beds, formations, provincial biozones, or, merely, glacial and interglacial deposits (Yakovlev, 1956; Flint, 1957; Zeuner, 1959; among others). It was only in the late 1950s (Brouwer, 1957; Leighton, 1958; Richmond, 1959; Van der Vlerk, 1955, 1957; Zubakov and Krasnov, 1959; among others) when the principles of the Pleistocene climatostratigraphic classification were first discussed owing to the construction of first National Codes of stratigraphic nomenclature. However, such terms as climatostratigraphy, climatostratigraphic subdivisions, kryomer, thermomer, climatolite and the like were introduced even later (Rozycki, 1964; Liittig, 1964, 1965, 1969; Zubakov, 1961, 1963, 1968; Krasnov, 1973; among others). In the USSR, climatostratigraphy owns its development to a theoretical dispute between its advocates and opponents, respectively, between the proponents of S.A. Yakovlev's school (Ganeshin et al., 1961; Zubakov, 1961, 1968; Zubakov and Krasnov, 1963; Krasnov, 1974) and those who adhered to the biostratigraphic school headed by V.I. Gromov; the latter proposed to use mammalian evolution as a crucial criterion for the subdivision of the Pleistocene (Gromov, 1948; Gromov et al., 1965, 1969). As a result, the principles and role of climatostratigraphy were accepted by the Soviet researchers (Menner, 1965, 1977, 1984; Shantser et al., 1973), including its former opponents (Nikiforova et al., 1982). However, we are still not in possession of a fully satisfactory climatostratigraphic theory or classification of the Pleistocene. Therefore, paleoclimatic studies are fraught with additional difficulties, when the procedures, techniques, and even terminology used in a certain climatostratigraphic correlation are to be specified. Hence, at present one cannot systematize data, differing in value and volume, on the past global climatic events without proper climatostratigraphic terminology. Climatostratigraphic criteria were first used for the subdivision of Pleistocene deposits in the glaciated area on the plains of Europe and North America. The glacial assemblages are complex in structure there: they are repeatedly interdigitated with lenses of fluvial and lacustrine deposits including lenses of buried peat bogs, while in the interfluves they are split by buried soils and weathering crusts. Intertill members yield plant remains and fossils that lived under conditions of temperate and warmer climate, as compared with the present. In other words, both the lithology and fossil content of the area suggest extremely drastic climatic changes that gave rise to repeated displacements of soil and vegetation zones by 15 - 20" in latitude and sometimes led to the entire disappearance of a forest zone. As early as the beginning of this century the glacial assemblages of the plains were found to have imbricate structure, that is glacial tills generally plunge to the north under the younger sediments and crop out in the southern part of the terraine where they form marginal belts of terminal moraines. As a result, the stratigraphic scale of glacial Pleistocene deposits was constructed using a morphostratigraphic principle proposed by Penck and Briickner (1909) to subdivide glacial events occuring in the Alps (Fig. 2.1) and later applied to the events occuring over the plains of Europe and North America. In so doing, it was assumed that the southern marginal belt of hummocky glacial topography was formed during the oldest cycles of continental glaciation known as Elsterian - Cracow - Oka glacials in Europe, Nebraskan and

41

Kansan stages in America, each subsequent glaciation being less extensive. All in all, there were recognized three or four glacial events; Table 2.1 shows correlation accepted in the 1940s. In the 1930- 1940s scientists (Eberl, 1930; Shapley, 1953; Yakovlev, 1956; Zeuner, 1959; Solar variations . . ., 1961; among others) could date glacial events only by comparing geological evidence with Milankovitch’s astronomical curve (1930) whose insolation minima were identified with glaciations on the basis of approximate estimates of the time interval during which lacustrine interglacial strata were accumulated. Although many workers (Markov et al., 1965; Flint, 1957, 1971) did not accept the hypothesis which relates glaciations to orbital disturbances, the Pleistocene stratigraphic scale remained unchanged, and the duration of the Pleistocene was either decreased to 300-400 ka (Markov et al., 1965) or increased to 1.2- 1.8 Ma (Flint, 1971). Borehole and palynological data which were widely used in glacial studies since the 1950s, have crucially changed our contention about the stratigraphy and chronology of glacial deposits and events occuring in the Northern hemisphere. The studies showed that the section is mosaic and incomplete: it hardly covers one-tenth of Pleistocene geological time, and there are no means of correlation between sequences; earlier correlation charts based on the morphostratigraphic principle, as a rule, proved to be incorrect and invalid (Van der Vlerk, 1955; Moskvitin, 1970; Kukla, 1977; Bowen, 1978). Therefore in the 1960s the global morphostratigraphic correlations “by main force” gave way to the recognition of local climatostratigraphic units based on the designation of stratotypes and on a comprehensive study

PTEROCARYA

w

I

0 L

L

an a

m

J

ww

0z

KJ

d a z

mm m -7- > c

’

v)

w1 w2

AA

I

DG

I

GM

I

MR

$-

i

I

RW

Fig. 2.1. Schematic profile to the gravel terraces in the Alpine foothills. The four youngest terraces represent the four glacial stages: Wiirm (W), Riss (R), Mindel (M), and Ciinz ( G ) of Penck and Bruckner (1909). “D” stands for Donau from Schafer (1953). RW, M R , GM and DO are erosional steps between the successive terraces, believed to represent interglacials. Ferreto is a red weathering zone on the upper terrace (cross-hatched). Soils and flood loams with interglacial fossils in black, loess stippled. M1, M2, R1 and so on show glacial moraines interdigitated with gravels. Note the deep erosion in Praeriss (PR). (After Kukla. 1977, fig. 18 and Kukla in Berger, 1981).

42

of interglacial deposits (Van der Vlerk, 1955, 1957; Richmond, 1959, 1970; Karlstrom, 1961, 1964; Luttig, 1964, 1965a; Zubakov, 1963, 1968b). Progress in climatostratigraphy was achieved mainly through deep-sea studies carried out in the post-war period when such new approaches as oxygen-isotopic curve (Emiliani, 1955, 1961, 1964), calcium carbonate content record (Arrhenius, 1952), and micropaleontological studies (Ericson, 1961; Ericson et al., 1964, 1968; among others) came into being. These pioneering works have shown that both deepsea sedimentation and ice-rafting reflect a global rhythmic pattern of climatic processes. 14C and U-series determination of the upper parts of deep-sea cores and interpolation of the obtained accumulation rates down the section promote the development of two competing chronological scales for the Pleistocene: short- and long-term time scales with ages of 350 ka for the glacial Pleistocene (Emiliani, 1961, 1964; Rosholt et al., 1961), and 2 Ma for the Gunz-Nebraskan (Ericson, 1961; Ericson et al., 1964), respectively. The second stage in the development of climatostratigraphy for the deep-sea Pleistocene started in the early 1970s with the introduction of the paleomagnetic method (Ninkovich et al., 1966; Hayes et al., 1969; Shackleton and Opdyke, 1973) and the factor quantitative analysis of microfossil assemblages which allowed the recognition of ecologically different water masses marked by certain limits of temperature and salinity (Imbrie and Kipp, 1971; Imbrie et al., 1973; Ruddiman and Mclntyre, 1976, 1981, 1984; Barash, 1983; among others). Apart from climatostratigraphy, which enables to ascertain and classify past climatic effects, it is very important to obtain numerical age estimates for the deposits. Correlation of climatic effects recorded in the deep-sea section and on land is possible only on the basis of reliable chronometric data. At present, Pleistocene deposits can be dated by a dozen methods, discussed in “Geokhronologiya SSSR”, Table 2.1. Traditional correlation of the Pleistocene glacial events in the Alps, Europe, and North America (G = Glaciation, 1 = Interglacial, Age in units of 10 ka) The Alps (Zeuner, 1959)

Wiirm (25-72-115) Riss-Wiirm

Northern Europe (Woldstedt, 1954)

Eastern Europe (Markov et al., 1965)

North America (Flint, 1957)

G

Weichsel

G

Valdai

G

Wisconsin

G

I

Eem

I Warthe

I

Illinoian

G

Holstein

I G I G I

Sangamon

Drenthe I

Mikulino Moscow Odintsovo Dnieper Likhvin

Yarmouth

I

Elster

G

Oka

G

Kansan

G

Afton

I

Nebraskan

G

11 (187)

Riss G

Saale G 1 (230)

Mindel - Riss

I I1 (435)

Mindel G

I (476) Giinz - Mindel

I I1 (550)

Giinz G

I (590)

Preglacial

Preglacial

43

volume 3 (Zubakov, 1974), and in a number of handbooks (Kaplin, 1976; Bowen, 1979; Punning and Raukas, 1983; Mahaney, 1984; Bradley, 1985; among others). Unfortunately, there is no reliable method for dating Pleistocene deposits, so the problem of the validity of obtained numerical estimates remains very acute. This is true not only for such experimental methods as thermoluminescence (TL), electron spin resonance (ESR), amino-acid racemization (D/L) and the like, but also for the theoretically faultless 14C and U-series methods. The stumbling block for the latter is isotope exchange between the samples to be dated and the enclosing rocks, because it leads to underestimated values. Therefore, at present only the crosscontrol of several independent methods can be used as a criterion of the validity of numerical data. The development of a Pleistocene chronostratigraphic scale amounts to a combin-

V28-239 tt

Vt6-205 0 -I

128

240 330 400

570

710

,900

Fig. 2.2. Key oxygen-isotope records. Formal numbering of 6 " 0 stages in arabic numerals after Emiliana (1961, 1964, 1978) and after Shackleton and Opdyke (1973, 1976), informal below stage 23 after Van Donk (1976); Terminations in Roman numerals after Broecker and Van Donk (1970); major cycles in capital letters after Ruddiman and McIntyre (1976). Warm oxygen-isotope stages stippled. Normal polarity in black, reversed blank. Y S - faunal zones of Ericson defined by the lack of the Globorotalia menardii tumida group. Faunal and floral markers showing: (a) beginning dominance of Erniliania huxkyi over Gephyrocupsa caribbeanica in transitional waters; (b) first appearance datum (FAD) of coccolith Emiliania hui-/eyi; (c) last appearance datum (LAD) of radiolaria Stylutracrus uniuersus; (d) LAD of coccolith fseudoemilianiu lacunosa; (e) appearance of abundant foraminifer Sphaeroidinella dehistens. (After Kukla, 1977, fig. 2). ~

44

ed usage of all the methods which are able to date climatic signals in broad frequency bands from hundreds of years to a few hundreds of thousands of years. In this case climatic signals are used to divide the deposits, while paleontological and chronometric methods help to “identify” and globally trace the climatic signals. Experience gained during the last decade showed that distant climatostratigraphic correlation is simpler with marine communities and taxa than with terrestrial ones, and it is probably simpler with loess rather than glacial assemblages. It became apparent that the stratigraphic division of the latter is complicated not only by the well known mosaic structure and lithofacies diversity of the section, but also by the pronounced glaciotectonics and, in particular, by numerous erratics of interglacial deposits. Hence, if 30 years ago the glaciated area was considered as a type area for Table 2 . 2 . Zonal biostratigraphic subdivision of the Pleistocene (age in ka) ~

Planktic roraminife, (Blow, 196’))

Globigerina calida calida Sphaerordinella subdehistens excawa N 23

Nannoplankton (Martini, 1976)

Homo saptens

Emilranla huxleyi N N 21

~

Gephyrocapsa oceanrca N N 20

_ _ _ (460) ~

lacunosa

(42-

loo?)

Manrmirrhrrs prirnrgenirrs - Microlus - Arvicola Dicrosronyx simplrcror

~

( 5 0 0 ) ~-

Pseudo emrlranla

Globororalia /run catulinoides N 22

Homo

Domestic fauna Dicrosronyl. rorquarus

- (220-270)

i’

~

~

Mammalian (Aleksandrova. 1976; Horai-ek, 19x1)

-

Crenolrrus doronicordes N N 19

(900)

Homo ererriis

M. lrogontherli Logurus rransiens - Prolagurus posrerius P1ioin.w

-

(Tiraspolian - Biharian Calerian)

~

Small

-~

( I 000) -

Gephyrocapsa

&(ll50)Helicoponrosphaera sellr

-(1800-2000)

-

(1500)-

A rchidiskodon meridionalrs lamanensis - Prolagurus pannonicus Prlymys hrntonr

-(

1200)-

Homo afarmsrs

45

the development of Pleistocene stratigraphy and deep-sea cores were tied in to the glaciated area, not it is quite the reverse - the deep-sea Pleistocene section (Fig. 2.2) is more and more often suggested be used as a stratotype (Kukla, 1977; Bowen, 1978; Berggren et al., 1980).

2.2. The significance of the oxygen-isotope scale for climatostratigraphic reconstructions At best, biostratigraphic methods can be used to divide Pleistocene deep-sea sediments into two or three zones (Table 2.2) which have no global distribution. Climatostratigraphy provides a more minute subdivision and more accurate correlation. It invokes the recognition of climatosedimentary cycles by three methods, namely micropaleontological, lithornineralogical and isotopic. The climato - micropaleontological method was proposed by Ericson and Wollin (Ericson, 1961, 1968; among others) who ascertained climatic zones using statistical estimates of the abundance of tropical species of planktonic foraminifera Globorotalia menardi (Orb.) - the most sensitive to temperature variations through the section. For the same purpose other workers used variations in the population of G. inflata (Orb.), Orbulina universa (Orb.), sinistral and dextral forms of Globorotalia truncatulinoides (Orb.) and Neogloboquadrina pachyderma and the like, along with cold-to-warm-water species percentage ratios (Parker, 1958; Barash, 1970, 1983; Ruddiman, 1971). Imbrie (Imbrie and Kipp, 1971; Imbrie et al.,

WBF

I1

Fig. 2.3. Comparison of 8 ' * 0 isotope curve and the WBF (warm-benthic foraminifera) curve based on cores from DSDP Site 397. The isotope curve was taken from Shackleton and Cita (1979). The WBF curve is based on benthic but different samples. (After Lutze, 1979).

46

I

I

+ 0 . 5 -0.5

-1.5

8'80 %o

1

-2.5

19 0

21 5 24.0 26.5 T,

OC

Fig. 2.4. Oxygen-isotope values (left) and winter sea-surface paleotemperature estimates (right) based on Caribbean core V 12 - 122. The sea-surface temperature estimates are derived from transfer functions. (After lmbrie et al., 1973).

1973; Sancetta et al., 1973) contributed greatly to this method. Using factorregression analysis he proposed to calibrate percentage ratios of present, ecologically representative, species with quantitative climatic parameters, such as winter and summer temperatures, salinity and the like, and to calculate these parameters down the section with the aid of transfer functions. Imbrie's method has found extensive application (Briskin and Berggren, 1975; CLIMAP members, 1976, 1986; Barash et al., 1983; Blyum, 1982; among others) and allowed as many as 10 to 15 units to be recognized in the Pleistocene section (Figs. 2.3 and 2.4). Carbonate cycles established by Arrhenius (1952) are widely used as well. The empirical data suggest that in the Pacific Ocean the CaC03 content is lower in interglacial deposits and higher in glacial assemblages. In general, the reverse is observed in the Caribbean Sea and in the Atlantic Ocean (Arrhenius, 1952; Hays et al., 1969; Gardner, 1982). In the Pacific this relation seems to be controlled by a higher rate of water mixing, while in the Atlantic it could be explained by a drastic increase in erosion of terrigenous material in glacial times. Core RC 11 - 209, one of the first paleomagnetically dated cores, is a stratotype for the recognition of carbonate cycles. In core RC 11 - 209, in the Brunhes orthomagnethem Hays and coworkers (1969) recognized nine Ca C03 content minima odd-number from top to bottom B (Brunhes) 1, B3, B5 and so on; in the Matuyama magnethem CaC03 minima are marked M1, M3 and so on; C a C 0 3 maxima are even-numbered B2, B4 and so on. Gardner (1982), who ascertained carbonate cycles in DSDP Holes 502 and 503 (Caribbean and Pacific areas, respectively) within the time interval up to 7.8 Ma proposed to number cycles, but not C a C 0 3 minima and maxima. For ex-

47

ample, he established seven cycles (Bl, 2 . . , 7) and 22 cycles ( M l , 2 . . . 22) for the Brunhes and Matuyama orthomagnethems, respectively (Fig. 2.5). The carbonate cycles differ in duration. For example, Gardner (1982) defined three ranks of cycles with amplitudes of about 40- 50 ka (C-cycles), 90 - 110 ka (super-cycles) and 500 - 600 ka. He considered super-cycles as a main pacemaker of the Ice ages. For core M 13519 through the western coast off Africa where the easterly winds permanently bear eolian dust from the Sahara, Sarnthein and co-workers (1984) constructed an age scale for carbonate cycles (CARPOR) assuming a constant accumulation rate of eolian material during the time interval between the last interglacial and the Brunhes - Matuyama boundary (Fig. 2.6). The variations of stable oxygen and carbon isotopes are the most important parameters in the stratigraphic technique of the Pleistocene. The variations are different in scale. Emiliani (1955) proposed 160/180 ratio shifts from the present-day level with an amplitude of about 1 Yo0 for isotopic half-cycles as corresponding, in his opinion, to Pleistocene glacial and interglacial stages. He proposed stages marked by lighter and heavier I 8 0 isotopes be odd- and even-number, respectively. Emiliani studied the isotopically lightest tropical planktonic foraminifera, such as Globigerinoides ruber, G. sacculifer and the like, whose tests are in isotopic equilibrium with ocean water they inhabit. Benthonic forms were studied at a later time; tests of some of them, e.g. Uvigerina sp., Planulina wuellerstorfii, Nonion sp., Cibicides sp., are also formed in isotopic equilibrium with ocean water. The life

7

Fig. 2.5. Correlation between carbonate stratigraphy of DSDP Site 502 and oxygen-isotope stratigraphy 0 1 V 28 - 239 and between the carbonate stratigraphies of DSDP Site 503 and RC I I 209 and oxygenisotope stratigraphy of V 28 - 239. The oxygen-isotope \[ages are modified from Shackleton and Opdyke (1976) and the carbonate cycles are from Hays et al. (1969). (After Gardner, 1982, fig. 10). ~

48 B R U N H E S

MAT.jJ.

u

_. 0

.

,

1

. 2

3

4

5

6

7

0

9

10

DEPTH ( R )

Fig. 2.6. Oxygen- and carbon-isotope records of benthic (Cibicides wuellersforJ0 and planktic (Globigerinoides sacculifer) foraminifera in core M 13519. Diamonds mark layers with enhanced bioturbation. Oxygen-isotope stages 1-21 indicated at the left margin. (After Sarnthein et al., 1984, fig. 2).

cycle of foraminifera being 30 to 40 days, the record of changing ocean oxygen isotopic composition at a specific level shows seasonal water temperatures. It is natural that curves constructed from studies of different species do not coincide. The bulk sample determinations yield averaged annual water temperatures. Both measurement techniques are in use. Interpretation of oxygen-isotope curves is strongly debatable. Assuming that the average isotopic composition of stored ice should be close to that of snow ( - 15%), Emiliani (1955, 1964) attributed 70% of the 6 I8O isotope shift to variations in water temperature. In this case, the water temperature in the Caribbean Sea must have been 6 - 8°C warmer in interglacial times. However, the record of changing isotopic composition of cores through ice sheets showed that they are composed of isotopically lighter ice ranging from - 20%0 at the margin of the Greenland ice sheet to - 50%0 in the Antarctic inland ice. As a result, later only 30% of the isotope shift varying from 1.6% to 2.2Yi during the glacialhterglacial cycle was attributed to a temperature factor (Dansgaard et al., 1969, 1971; among others). The remaining 70% of the shift was attributed to changes in ocean oxygen isotopic composition

49

due to redistribution of l80and l6O between the ocean and continental ice at different stages of the glacial/interglacial cycle. Independent determinations of temperature changes in the surface water layer during the two last glacial/interglacial cycles obtained by the micropaleontological method yield the value 2-3"C (CLIMAP members, 1984) which is adequate for an isotope shift of - 0.5%0. Therefore the majority of workers believe that the oxygen-isotope curve implies primarily global changes in the volume of ground ice and ice shelves, i.e. it should be regarded as a paleoglacial curve. Core V28 - 238, the first to be analyzed both isotopically and paleomagnetically, was suggested be used as a stratotype for oxygen-isotope cycles (Shackleton and Opdyke, 1973). At present a dozen cores are subdivided in great detail into isotopic cycles and dated at the base of the Brunhes - Matuyama boundary (Table 2.3). In all the cores the reversal corresponds to isotope stage 19 (Fig. 2.7). We can determine the age of isotopic cycles through extrapolation of accumulation rates. Independent radiometric dates were available now for some isotopic stages. For example, a stage 11/12 boundary date around 440 k 40 ka was obtained from K - Ar analysis in ash; the age of stage 12 was estimated at 420-460 ka by Saito using the ESR method and around 474 ka by Wintle and Hantle using the TL procedure (Sarnthein et al., 1984). Emiliani (1955, 1964, 1978), Broecker and Van Donk (1970), Hays et al. (1976), Berger et al. (1981) and others have found a morphological similarity between the isotope variation curve (Fig. 2.8) and the solar insolation curve plotted by

Table 2.3. Key sites of oxygen-isotope stratigraphy with Brunhes - Matuyarna transition ~-

~~

4s

Latitude and longitude

Reference\

~-

P 6304

~

9

V 28 - 238

14"57'N, 6R"SS'W

'N,l60"2Y 'E

2. I I.65

Glohiyerinoides racculiJer (pla) GlohiRermoides

~

670

Emiliani (1966)

880

21'18'N. 22"41'W

1.2

Glohryerinoirlr~ wcculr./er (pla)

> 730

15Y"II'E

I.o

> 2000

14"55'N. 69"OS'W

2.2

Clohigerinoides socculijer (pla) ClohiRerinordes sacculifer (plal

780

Shackleton and Opdykc (19731 Parkin dnd Shackleion (1980) Shackleton and Opdlke (1976) Emiliani (1978)

25"30'S. I I " 1 8 ' E

2.2

UI ixerino sp.

> 730

Morley and Ha\,

1'01

mcculifer (plal

v 23

v

~

100

28-239

P 6408

~

9

KC 13-229

3'15".

( b e ) Globoro-

(1981)

ralia inflora

(pla) Cl 984 DSDP 552A

V 22-174 DSDP 504 DSDP 502B .vl 13519

56"02, 56".

20

Planktic species Benthic 5peciss

23"13. 3 8 ' W IO"O4'S. I 2 " 4 Y ' U 1.14". 83'44'W

I .Y

15.Jacculifer

> 730

4.4

>

ison

Il"30'N. 79*23'W 5"40'N. I Y " 5 1 ' W

2.25 14

G. AocculiJer G. rirber (pla) G xm-u/iJer Ci w c u / t f e r Phnulino wueller

-

1700 775

48"S, 135"W

srorfii

(be1

970 ~

3500

NikolaeL (1981) Shackleton er al. (1982. 19841 Shackleton ( I Y 8 2 ) Shackleton and Hall (1983) Piell (1983) Sarnthein ct nl. (1984)

50

DSDP 5 0 2 B 6 "0

(%o)

PDB

0.50 0.00 -0.50 -1.00 -1.50

0 2 4 c)

6

N

0

-

I

0

1

8

I +

a

10

c2

12

w

5

0

14 0

m 1 m

16

3

18 D

w

;2 0 W

g

22

0 V

24 26 28

0

502 B 502 A

30 0,

32

N

0

34

0

36

0 N

0 ln

38 4c

0

Fig. 2.7. The oxygen isotopic records for hydraulically piston cored DSDP Site 502 B from the Caribbean Sea based on monospecific analyses of Globigerinoides sacculifer. The amplitude of 6I8O change increases and its mean value becomes more positive after 900 ka BP. (After Prell, 1982).

51

Fig. 2.8. Long-term variations of climate over the previous 400 ka. The continuous line represents the simulated climatic variations obtained from the regression model. (From Berger et al., 1981). 1 insolation variations corresponding to asiranomical model of Milankovitch (1930); 2 - 6’’O values for cores RC 1 1 - 120 and E 49- 18 (after Hays et al., 1976); 3 - climatic curve (after Berger et al., 1981). ~

0.4 Age, Ma

06

Fig. 2.9. Oxygen-isotope record derived from DSDP Site SO4 for the past 700 ka, based on analyses of C. saccuh’fer compared with record derived from conventional piston core V 28 - 238, both plotted to a uniform accumulation. (After Shackleton and Hall, 1983).

Milankovitch (1930) and then confirmed and specified by Sharaf and Budnikova (1967), Sharaf (1974), Vernekar (1972), and A. Berger (1978, 1979). Spectral analysis of some oxygen-isotope curves enabled periods around 12 - 23, 41, 96, 250 and 413 ka to be recognized therein, caused by changes in the geometry of the earth’s orbit: changes in eccentricity, obliquity and precession (Hays et al., 1 969, 1976; Komintz et al., 1979; Briskin and Harell, 1978; Morley and Hays, 1981). As

52 Table 2.4. Age estimates (in ka) for oxygen-isotope boundaries Isotope stages

Shackleton Emiliani (1978); Berggren et al. (1979) and Opdyke (1973)

Komintz et al. (1979)

Morley . and Hays

26AI

TWEAQ

(1981)

Age accepted in this study

.~

Rate of sedimentation 4

75

72

80.4

73

72

73

5

128

-

138

127

128

128

195

163

209

190

188

190

25 1

-

266

247

244

245

297

259

307

216

279

285

347

-

348

336

334

335

367

337

370

352

347

350

6 7

8 9 10

11

436 453 42 1 (K/Ar - 440, TL 474, ESR 420 - 460) 465 480 475

440 12

425

472

404

502

-

502

510

505

505

542

466

540

55 1

517

525

592

-

585

619

519

585

627

5 39

619

649

608

600

640

662

67 1

660

480

13 14

I5 16

17 647 18 19 N , / R ,

688

632

...(690).. .

...(690)...

...(690).. .

677

712

724

700

. ..( 693)...

...(728) ...

...(730) ...

...( 734) ...

693

750

744

745

700 20 (726)

718

726

760

756 - 776

830

21 (756) 22 23

809 - 830 (809) (Upper Jaramillo reversion, K/Ar 900)

890 920

24 940 25

(Lower Jaramillo reversion, K/Ar 970)

53 6 “0 %.

Isotope stages

5

7

9 2.5

If

fJ

6

17

83

6.I

fE

23

21

25

27 25

f2.9

Depth,m

Fig. 2.10. Curve A. Isotopic stratigraphy provided by b I 8 0 variations of benthic foraminifera Cibicides, Planulina and others, normalized to Uvigerina. Curve B. Globorotalia truncatulinoides coiling ratios, southward excursions of the Brazil current might be indicated by percentages of right-coiling forms higher than present, 40% (shaded area). Curve C. Downcore 6 ’ * 0 record of Globigerinoides ruber at Hole 517 with time. (From Vergnaud-Grarzini et al., 1983, fig. 6).

a result, there were constructed several time scales for isotopic cycles whose boundaries were corrected using orbital parameters (Table 2.4). New oxygen-isotope curves for DSDP Hole 504 (Shackleton and Hall, 1984) and DSDP Hole 517 (Vergnaud-Grazzini et al., 1983) cores and for core M 13519 (Sarnthein et al., 1984) showed even better agreement with astronomical parameters than than for core V 28 - 238. Figs. 2.6,2.9 and 2.10 show that the new curves have more clearly defined thermomeric isotopic peaks for stages 9, 11, and 21, along with kryomeric3 peaks for stages 16 and 22. The amplitude of the isotope shift is also greater, particularly for Terminations IV, V , VII, and X, attaining 2.29‘00 and even 2.49‘00in core M 13519 and DSDP Hole 517 core, respectively. This requires a new interpretation of the isotopic data and suggests that the isotopic of ice composition varied radically during stage 22 (Vergnaud-Grazzini et al., 1983). In recent years, particular emphasis is placed upon the 12C/13C isotope ratios. The light isotope I2C is readily accumulated in planktonic forms and carried onto the ocean floor within the shells of dead organisms. Therefore, variations in the 6 I3C curve are opposite to those in 6 I8O curve (Fig. 2.6). Nevertheless, they show a close similarity which implies that variations in both 6 13C and 6 ‘*O are affected by temperature and salinity changes. Moreover, variations in 6 13C reflect changes in the productivity of sea surface waters and the rate of vertical circulation, i.e. changes in the strength of bottom currents and upwellings (Kroopnick et al., 1977; Duplessy, 1981). Although the mechanism and causes of the variations in 6 I3C are not fully understood, in general they may be probably controlled by the carbon dioxide redistribution between the atmosphere, the ocean surface layer, the deep-sea water, and the terrestrial biota. Recently Shackleton and co-workers (1983) showed that variations in the 6 I3C As the word is derived from the Greek “kr);os”, the author, following Luttig (1964). retains the spelling with initial k.

54

curve really reflect changes in the carbon dioxide level in the atmosphere which were measured for the last 40 ka in air bubbles from cores through the Greenland ice sheet (Delmas et al., 1980; Neftel et al., 1982; Oeschger et al., 1985). Comparison of the oxygen-isotope curve, adjusted to CO,, to that of solar insolation for the last 340 ka led Shackleton and Pisias (1984) to the conclusion that changes in insolation caused by orbital parameters precede variations in 6 13C and in CO, levels in the atmosphere, which, in turn, occur prior to climatic fluctuations, i.e. changes in ice volume and temperature (for details see Summary). Hence, all the studied cycles, namely, micropaleontological, carbonate, oxygenisotopic and carbon-isotopic, seem to be correlative with each other, with variation of the CO, level in the atmosphere, as well as with changes in solar insolation. Consequently, all of them represent different facets of a single complex process, namely variability of the climatic system which seems t o incorporate the atmosphere, the ocean, the kryosphere, the biosphere with due regard for astronomical factors. The foregoing suggests that the above listed cycles should not be regarded as independent stratigraphic subdivisions. They are only certain aspects of a single climatostratigraphic zonation. However, in fact oxygen-isotope cycles ought to be given preference, primarily because they can be established both in deep-sea and continental sections. On land the record of changing oxygen isotopic composition can be investigated in lacustrine carbonates, travertines, and especially in ice. Cores through ice sheets provide important information on climatic changes for the last 100- 150 ka (Dansgaard et al., 1971, 1982, 1984; Oeschger et al., 1985), as shown below (Fig. 2.11). In order to unify climatostratigraphic terms, the authors propose to introduce a term “orthoclimathem” (OCT) to define: (i) isotopic stages of Shackleton and Opdyke (1973); (ii) carbonate half-cycles of the scales proposed by Hays and coworkers (1969) and Gardner (1982); and (iii) appropriate climato-micropaleontological zones. This term is used for global climatosedimentary cycles, regardless of the technique used t o ascertain them, with approximate duration measuring tens of thousands of years. Stratotype core V 28 - 238 for isotopic stages (Shackleton and Opdyke, 1973) is suggested be used as the orthoclimathem stratotype and core M 13519 is proposed be taken as the regional Atlantic parastratotype (Figs. 2.2, 2.6). Micropaleontological datum levels related to the first and last appearance of taxa or morphological forms over the section (FAD and LAD, respectively) are essential for interregional correlation of orthoclimathems in the marine sections (see Fig. 2.12).

2.3. Systematic aspects of “ocean - continent” climatochronological correlation. The significance of geomagnetic data In one of his earlier works the author (Zubakov, 1968c) wrote that the development of chronological scales for the deep-sea and continental Pleistocene should be independent and each scale should be based on its own empirical data. However,

55

recent advances in deep-sea climatostratigraphy and, primarily, the development of the oxygen-isotope curve have put a different emphasis on the role of oceans and continents. Kukla (1977) was right in his statement about the scale of oxygen-isotope stages as a real standard for the unified Pleistocene time scale. He was also the first to attempt to relate climatic events occuring in Europe to isotopic stages (Fig. 2.2). Similar attempts were made later (Bowen, 1978,;Lindner, 1980; Voznyachuk, 1978, 1985; Nikiforova et al., 1982; Liu Ze Chun, 1982; Wiegank, 1982; Bonifay, 1983; Zubakov and Borzenkova, 1983; among others) in a variety of versions. But ambiguity of the contentions suggests a lack of valid data for such a correlation. I n most cases correlation was based on the number of climatic cycles recognized above the Brunhes- Matuyama boundary. At best, TL datings were used. However, their accuracy is only k 2 5 % and, hence, they cannot be relied upon. Moreover, a number of important questions that can be raised about correlation between con-

0,

'

-35 "

;3,5, , , -,310L:/o,

-30 "

' I

" I

,

10 -

20

-

30 40 -

50 60 70 -

80 90 -

700-

110-

120ko

B.P. -

Fig. 2.1 I . Profiles of 6 ' * 0 measured in Copenhagen along the Dye 3 (0 to 1982 m deprh) and the Camp Century (0 t o 1370 m depth) ice cores plotted on a common linear time scale based on considerarionr discussed by Dansgaard et al. (1982). In the time interval 40-30 ka BP on the left side of the Dye 3 b"O profile, the core increments analyzed in more detail regarding CO, concentrations in Fig. 10.5 are indicated. (After Oeschger et al., 1985, fig. 4).

56

tinental and deep-sea events have no clear answers. These questions are as follows. Why is the number of kryomers in the Pleistocene isotope scale larger than the number of traditional glaciations? Why is the duration of interglacial stages on land estimated around 10- 15 ka, while the duration of kryo- and thermomers of the isotope scale is equivalent? Why is the value of observed glacio-eustatic sea level variations - 100- 120 m - almost half that calculated from isotopic data (6 lSO glacial/interglacial shift averages 2 f 0.4%0 which approximately corresponds to the volume of ground ice equivalent to a 190-210 m ocean water layer)? 12 3

5

6

7

8

9

22ll

11

12

13

14

15 16 17 16 19

v34-54

4.2 J v34-53

227

V34-52

$

-0.2 0.3 0.8

J

0

100

200

300

400

500

600

700

ROO

Time ( k y r s a g o )

Fig. 2.12. Variations in 6 ' * 0 as a function of estimated age for six of the transect cores. Key biohorizons (coiling direction change in G. rrussuformis, local last appearance of pink G. ruber, P. lucunosu LAD), and the stratigraphic level of the Toba Ash are identified. (After Peterson and Prell, 1985, fig. 3).

Nikolaev (1981, 1984) proposed an absolutely new interpretation of the oxygenisotope scale. He claimed that measurements of isotopic composition of planktonic foraminifera G . ruber, G. sacculifer and the like living in surface waters yield false “light” maxima associated with the periodically repeated surges of ice into the ocean in volumes which might have accounted for one third of the present-day ice volume in Antarctica. In Nikolaev’s opinion, analysis of bulk samples of the entire foraminifer assemblages, along with independent climatomicropaleontological data (Blyum, 1982), suggests that the past 700 ka witnessed only four or five coolings. In particular, isotope stages 3, 7 , 13, and 23 may well be false thermomers. There are known many discrepancies bet ween the isotope and insolation curves, which were emphasized originally by Shmuratko (1984). Summary accounts of geochronological data on continents (Zubakov, 1968, tables 2 and 6) and on the territory of the USSR (see Zubakov, 1974, table 48), compiled prior to the introduction of Shackleton and Opdyke’s oxygen-isotope scale, also report no more than five or six glacials for the last 700 ka. Thus, only a valid correlation of radiometrically dated climatic events in the Pleistocene deep-sea and continental sequences can give an answer to the question of whether all the isotope stages correspond to independent glacials and interglacials. Correlation of deep-sea and continental sequences has proved to be one of the most difficult stratigraphic problems. Data on geomagnetic field fine structure, i.e. on changes in a field younger than 100 ka, are vital when the Pleistocene is concerned. The changes are divided into events, excursions and secular variations. An event is a short-term magnetic field polarity reversal relative to the adjacent intervals. An excursion means a sharp fluctuation of the field which does not result in complete reversal but causes a shift in direction of the virtual geomagnetic pole by more than 45” in latitude. Secular variations are smooth low-amplitude changes in direction of the field. Excursions are characterized by an anomalous state of the geomagnetic field, namely, by a sharp intensity drop and a large scatter in the direction of the I, vector reminiscent of reversals, i.e. they may be considered as “incomplete reversals”. On earlier scales (Khramov, 1958; Cox et al., 1964) the last polarity epoch Brunhes orthomagnethem - was interpreted as monopolar. However, as early as 1966 Ninkovich and co-workers reported the presence of four horizons of anomalous polarity in core V 20- 108, and Smith and Foster (1969) recorded the Blake event of reversed polarity with an age of 108 - 114 ka from six deep-sea cores. At the same time excursions of the geomagnetic field were recorded in the continental section as well (Bonhommet and Babkine, 1967). At present 8 to 12 sharp fluctuations of the geomagnetic field are distinguished in the Brunhes orthomagnethem (see table 2.5). However, their interpretation is ambiguous. Some authors (Harrison, 1974; Pospelova and Gnibidenko, 1982; Pisarevsky, 1983; Tretyak, 1983; among others) consider them as global, while others (Olausson and Svenonius, 1975; Pevzner, 1973, 1982; Thoveny et al., 1985; and others) believe that they are regional or even local fluctuations which cannot be dated precisely. In the author’s opinion, the geomagnetic datum marks are the only, aside from climatic ones, which are recorded in sediments derived from any source and can be

58

synchronous on a global scale. Thus, the magnetochronological scale of excursions and events is an independent “tool” for global correlation of the Pleistocene deposits. Parallel climato- and magnetostratigraphic subdivision of the sections ensures mutual control of the two methods. There is no sense in rejecting the debatable character of the scale of excursions. It has become the subject of study only in very recent years; of course, some excursions would prove to be false ones and some others would be reclassified as events. Therefore all data currently available on excursions and events, particularly over continuous sections, should be systematized. No less than ten complex geomagnetic events, apart from the BrunhesMatuyama, are recognized for the last 1.2 Ma. Events revealed in deep-sea sections are shown in histograms compiled with the aid of statistical techniques (Fig. 2.13). Table 2.5 presents geomagnetic datum marks for the Pleistocene with ages determined by radiometrical, paleontological or archeological methods. Four groups of data were used, namely deep-sea cores (Clark, 1970; Wollin et al., 1971, 1977; Ryan, 1972; among others), radiometrically dated lava flows (Cox et al., 1964; Champion et al., 1981; and others), a unique section at Lake Biwa in Japan (Jaskawa, 1974, Horie, 1982), and continental loess-soil series known in the USSR. The latter were well studied from six areas, namely from terraces of the Black Sea and the Sea of Azov (Kochegura and Zubakov, 1978; Tretyak, 1983; Vlasov et al., 1983, Puleomagnetic . . ., 1983), the Dniester Valley (Kulikova, 1980), Tissa Valley (Adamenko et al., 1981), the upper Ob River (Pospelova and Gnibidenko, 1982), and from the piedmont depressions of Soviet Central Asia (Penkov et al., 1976; Lazarenko et al., 1980; Dodonov and Ranov, 1984). The substantiation of climatostratigraphic position of four young excursions was obtained from diluvial shelves on the surface of terraces I1 above the flood-plain

60

70

00

ka

Fig. 2.13. Histograms showing the geomagnetic excursions revealed in deep-sea cores with further statistical processing of data: A - after Pisarevsky (1983), B - after Kochegura (in Zubakov, 1974). I - with 50 ka time interval; 2 - with 25 ka time interval. Formal numbering of excursions in arabic numerals after Zubakov (1984) - see Table 2.5 and Fig. 2.14.

59 Table 2.5. Geomagnetic datum planes in the Pleistone, estimated age in ka. .- .

Author’s Geographical nom- Type section, age nomenenclature, references clature

~-

. -

1

2

-

Reliable correlation

Position with respect to isotope scale - -

4

3

5 ~-

.~

rla

Etruscan

Etruscan ceramics, ca. 2.5

rlb

Gothenburg (Morner, 1971)

Glacial varves, Gothenburg, Sweden, I4C 12.3- 13

r2

Mono (Denham, 1974)

Lake Mono, California ‘‘C 24

~~~

1

-~

Lake Erie, I4C > 7.6< 14 (Creer et al., 1976)

~-

r3a

Laschamp (Eonhommet and Babkin, 1967)

1/ 2

~

~~~

Lake Mungo, “C 29.5 (Barbetti and McElhinny, (1976); Korman, Dniester River, 14C 23.5 -26.5 (Kulikova, 1979); Gmelin site, Kostyenki, Don River, “C > 22, TL < 26.5 (Pisarevsky, 1983)

2

Laschamp and Olby Volcanoes, Central French Massif, T L I4C 36 5 4 (Gillot et al., 1979) ~

~~

3 ._

r3b

Olby (Gillot et al., 1979)

Laschamp and Olby Volcanoes, Central French Massif, K/Ar 42 i 5 (Gillot et al., 1979)

Molodovo, Dniester River, I4C 43.3 -44.0 (Kulikova, 1979). Kargopolovo, O b 1.5 (PospeRiver, I4C 41 lova, 1981)

r4

Blake (Smith and Foster, 1969)

8 deep sea cores, ca. 108- 1 I4 (Smith and Foster, 1969), K/Ar 113 5 2 Creer and Readmen, 1980)

Brno (Kukla and Kochi, 1972) BI cycle; Kostyenki IV, Don River, TL < 170 k 40; and Chokrak 11, TL < 165 5 40 (this work)

Biwa I (Jaskawa, 1974)

Lake Biwa, Japan, F.t. 176-186 (Jaskawa, 1974)

13 Chegan I , Altai, T L 145 (Faustov et al., in Zuba-, kov, 1971, 1973)

Jamaica (Wollin et al., 1971; Ryan, 1972)

Core V 12- 122 (Wollin et al., 1971), c a . 200

Chokrak 111, Kerch Peninsula, T L 165 i 30 (Vlasov et al., 1983)

- .~

___

r5a

~~~~~

-

5d

-~

6/7

-.

r5b

6/7

60 Table 2.5. (continued) Author’s Geographical nom- Type section, age nomenenclature, references clature

Reliable correlation

Chegan 11, Altai, T L 266 f 30 (Faustov et al., in Zubakov, 1971, 1973); Tsokur 111, Taman Peninsula, TL 285 r 48 (Zubakov and Kochegura, 1971); Levantine ( = “Y”) core RC 9 - 181, ca. 280-320 (Wollin et al., 1971; Ryan, 1972)

Position with respect to isotope scale

Odintsovo - Galich (Trukhin, 1969)

Dnieper till, Odintsovo - Chekalin, TL 280 i 30 (Faustov et al., in Zubakov, 1973)

Biwa 111 (Jaskawa, 1974)

Lake Biwa, Japan ca. 350-367 (Jaskawa, 1974)

r7a

Emperor (Ryan, 1972)

Core RC 9 - 181, Mediterranean, ca. 390 - 400 (Wollin et al., 1971)

Tioga-Tauro 0,ca. 400 Creer and Readmen, 1980)

11

r7b

Ureki I (Zubakov and Kochegura, 1971, 1973)

Paleo-euxine beds, Ureki, Georgia, T L > 330, ca. 450 - 500 (Zubakov and Kochegura, 1975)

Snake River 11, K/Ar 460 f 50 (Champion et al., 1981)

12

r6a

r6b

8/9

1O?

~

r8a

r8b

Ureki 11 - Jakhno Tsokur-Yakhno, (Zubakov and Taman Peninsula, Kochegura, 1971) 8th soil (Zubakov and Kochegura, 1971 - 1973); Chokrak - Patrayi beds, T L 580 t 140

Chervony-KopeE, ca. 500 f 50 (Bucha, 1973); Elunino 1, Ob River (Pospelova et al., in Zubakov, 1971)

Mosty - “Oka” Byelorussian (Zubakov and ( = “Oka”) till, Kochegura, 1973, Mosty, Niemen 1974) River, TL 560 f 60 (Zubakov, 1974)

Tsvermagala Mountain, Georgia, Baku beds, TL 520 130 (Zubakov et al., 1986); “Don - Orlovka”, Don till (Krasnenkov et al., 1984) Elunino 11, Ob River (Pospelova et al., 1982)

~-

15

16

61 Table 2.5. (continued)

_ _ _ _ _ _ _ Author’s Geographical nom- Type section, age enclature, references

_ - _ _ _ Reliable correlation ~~

--- -

_-

R / N transition Cheleken PeninCore V 28 - 238, Pacific (Khramov, 1958) sula, Baku/Apsher(Shackleton and Opdyke, 1973) Matuyama - Brun- onian boundary hes reversion (Cox Khramov, 1958), et al., 1964) K/Ar 690 (Cox et al., 1964). K/Ar 730 (Mankinen and Dalrymple, 1979) Unnamed (Hirioka Deep-sea cores, ca. 830 (Watkins, and Kawai, 1967; Watkins, 1968) 1968), Japan, K/Ar 850 i 30 (Hirioka and Kawai, 1967)

Lake Zykh, Apsheron Peninsula, K/Ar 820 k 250 Kvemonatanebi I ? (Zubakov and Kochegura, I973 1974); Taylor Valley, K/Ar 840 f 30 (Mankinen et al., 1981)

Jaramillo (Cox et al., 1964)

Jaramillo, New Mexico K/Ar 850- 900 (COX, 1969); Clear L.ake, California, K/Ar 900- 970 (Mankinen et al., 1981)

Core V 28 - 239 (Shackleton and Opdyke, 1976)

Cobb Mountain

Clear Lake, Coso Range, California, K/Ar 1100 k 20 (Mankinen and Gromme, 1982)

Argentina, K/Ar 1050 (Fleek et al., 1972); Komioke, Japan F.t. 1100 (Maenaka, 1975 1979); Kvemonatanebi I I ? ca. 1100 (Zubakov et al., 1975)

Position with respect to i5otope scale .~-

19

22/23

-~

23/24

- .~

25/26?

~

_

_

in the Dniester River basin, in the area of Paleolithic Molodovo and Korman sites (Kulikova, 1980) and, particularly, in the Don River basin, near Kostyenki village (Pisarevsky, 1983). Coastal sections on the Kerch Strait shelf (Zubakov, 1973, 1974; Vlasov et al., 1983) and a section at the Lower Paleolithic Korolevo site in the Tissa River valley in the Transcarpathians (Adamenko et a]., 1981) show standard values for older excursions. These sequences contain a continuous succession of six to eight excursions for which thermoluminescence (TL) ages and archeological data are available; the ages are also supported by the relation between buried soils and marine terraces (Fig. 2.14). Analysis of Table 2.5 and materials used in its compilation reveals that: (i) terrestrial and marine sections in the Brunhes ort homagnethem contain approximately

62

?

\ \

63

the same number of excursions - about 12 (or 8 - 9 if some were considered as double); (ii) the stratigraphic positions of the excursions relative to buried soils and isotope stages in the terrestrial and marine sections, respectively are, in general, similar (although some “vagueness” of the excursions probably due to soil bioturbation and degradation may be noted); (iii) numerical values for excursion ages are not affected by the geographical position of a section. Taken together these factors suggest a global character of the excursions. Therefore, it is obvious that not only complete reversals of the geomagnetic field, but also “incomplete reversals”, i.e. excursions, may be used for minute interregional stratigraphic correlation, although only in combination with data obtained by other methods of dating, such as biostratigraphic and radiochronometric, enabling identification of reversals and excursions. Only in this context excursions, as well as complete reversals, may be considered as “traces” (indications) of age in years, but needless to say these ages should be considered only as provisional. Let us now direct our attention to Table 2.5 and consider the sequence of geomagnetic polarity events taken by the author as reliably established for the Pleistocene. (1) The Holocene geomagnetic polarity event (rl) includes Etruscan ( r l a ) and Gothenburg (rlb) excursions. The Etruscan excursion was identified from kilns used to bake amphoras; they were dug out in the sites of ancient Etruscan settlements in the central part of western Italy. The historical age of the excursion is about 2,600 yr. The Gothenburg excursion was recognized in 1966 by Ninkovich in core V 20 - 108 through the Northern Pacific, and by Morner (1971) in varved clays from the Botanical Gardens in Gothenburg, Sweden. Its 14C age is around 12.3 - 13 ka. (2) The Mono (Denham, 1974) - Mungo (r2) event was recognized in lacustrine deposits of California; its I4C datings are around 24 ka and 29.5 ka for Lake Mono and Lake Mungo, respectively. An event with I4C ages ranging from 23.5 to 26.5 ka was also established in the section of the Upper Paleolithic Korman site, Dniester Basin (Kulikova, 1980). Pisarevsky (1983) discovered the event dated by the TL method at 26.6 ka in the section at the Gmelin site, near the village of Kostyenki, Don River area. (3) The Olby - Laschamp ( = Kargopolovo) event was established by Bonhommet and Zahringer (1969). Later the event was divided into two excursions (Gillot et al., 1979), namely, the earlier - Olby - excursion (r3a) with K/Ar age at 42 k 5 ka was studied on lavas from the Central French Massif; the excursion with I4C age

Fig. 2.14. Geomagnetic excursion derived on the Pleistocene key sections of the Black Sea area, after Zubakov and Pisarevsky. Sections: I - Kostyenki, Don River; I1 Cape Tuzla, Taman Peninsula; 111 - Chokrak Lake, Kerch Peninsula; IV - summary column. D declination, I - inclination, 4 - latitude of virtual geomagnetic pole; K 1 4,T 1 - 6, C 1 - 5 local zones of anomalous polarity; r 1 - 8b - formal numbering of excursions established for regional nomenclature. excurLegend: 1 - loess; 2 - buried soil; 3 - lagoon silt; 4 - marine sand; 5 - normal polarity; 6 Fion; 7 - anomalous polarity. ~

~

~

~

64

around 43.4-44 ka was also recognized in the terrace 11 section, Ob River at Kargopolovo (Pospelova and Gnibidenko, 1982) and in the Dniester Basin, near Molodovo (Kulikova, 1980). According to Gillot, the TL and 14C ages of the later - Laschamp - excursion (r3b) is about 36 t- 4 ka. (4) The double Blake event (r4) was initially recognized in deep-sea cores (core A 179-4, Caribbean Sea, and some others) in the X-zone in sense of Ericson ( = isotope stage 5) by Smith and Foster (1969); earlier estimates showed an age range of 108- 114 ka, but recently Crear and Readman (1980) obtained a date of 113 k 2 ka from K/Ar analysis. The event is most often established in Pleistocene continental sections. (5) The Jamaica - Biwa double event (r5). Jamaica excursion (r5b?) (rebound of 70" inclimation) was determined by Wollin, Ericson and Ryan (1971) in core V 12- 122 through the Caribbean Sea, in isotope stage 7 (after Van Donk). There were also grounds to recognize a weaker anomaly of 40" inclination in stage 6; later it was established by Yaskawa (1974) in the Lake Biwa section, Japan, as Biwa I having a fission-track age range of 176- 186 ka (r5a?). Its equivalent in the Chegan section, Altai Mountains, is R-zone with TL age at 145 ? 13 ka, recognized by Faustov and co-workers (see Zubakov, 1973, 1974) in the Mayma tills. Both excursions with 80" and 60" inclination were recorded by Ryan in core RC 9 - 181 taken from the Mediterranean. (6) The Levantine excursion (r6) with 50 - 60" inclination maximum was identified by Ryan (1969- 1972) in core RC 9 - 181, in isotope stage 9 after Emiliani. According t o Steisi, it is Y-zone dated at 320 ka. Subsequently, Trukhin (Sudakova et al., 1974) recognized the R-zone, later named the Odintsovo - Galich, in varves from the Dnieper tills; a similar R-zone, Chegan, was established by Faustov and co-workers in 1971 (Zubakov, 1973) in lacustrine clays separating Mayma and Katun tills in the Chegan section. The K TL age of this zone is about 266 k 30 ka. At the same time Zubakov and Kochegura (1971) reported data on the double R zone Tsokur, covering Dnieper loess and underlying soil 5 at the Tsokur section, Taman Peninsula. Its K TL age is dated at 285 t 48 ka. The two-stage structure of the zone was supported by the data from Lake Biwa section where Biwa I 1 (r6a) and Biwa 111 (r6b?) excursions were dated at 292 - 298 ka and 350 - 367 ka, respectively (Yaskawa, 1974). (7) The complex Ureki event ( = Ureki I) - r7 - was ascertained by Zubakov and Kochegura (1973, 1976) in fossiliferous paleo-euxinic beds from coastal sections in Georgia; according to Shelkoplyas, the K TL age of the event is no less than 330-350 ka, but Zubakov and Kochegura believe that the event occurred 450-500 ka BP. Earlier Zubakov and Kochegura (1970) recorded the event in buried Tsokur soil 7. Bucha (1973) found its effects in the Cerveny - Kopec loess section; its age was estimated at 500 t 50 ka. Its probable equivaIent in the deepsea section may be the double Emperor excursion recognized as a rather weak, 30", anomaly of inclination in cores V 20- 108, V 12 - 122 and RC 9 - 181 from sediments assigned to isotope stages 11 and 12 with age estimated at 350-485 ka. This event was reliably confirmed by three K/Ar datings (515 f 85, 400 k 75 and 495 f 90 ka) of lava flows 10 and 1 1 from the lower part of the lava sequence in

the Snake Valley; the age of R-zone Snake 111 averages 460 - 475 ka (Champion et al., 1981).4 (8) A double excursion with an age of 600 ka (r8) was first recognized in deep-sea cores by Wollin and Ericson in 1971; in 1972 Zubakov and Kochegura found it in lower (Oka) loess of the Tsokur section at Yakhno farm, as well as in barren clays of the Ureki section beneath the beds containing paleo-euxinic fauna. The excursion was also recognized in ancient Dzukiya tills from a borehole drilled near the village of Mosty, Niemen River area; the age of the excursion determined by the K TL method is about 560 k 60 ka (Zubakov, 1974). The Mosty- Yakhno- Ureki 11 excursion corresponds to an anomalous, according to Toichiev, zone with K TL age at 520 t 130 ka in the Shava beds of the Tsvermagala Mountain (Zubakov et al., 1986). A two-stage character of the same (?) excursion was confirmed by Pospelova (1982) in the loess section near the village of Elunino, Ob River basin. (9) The double Jaramillo event (nZ) in the Matuyama orthomagnethern was first reported by Cox and co-workers (1964) from lava flows, New Mexico; its earlier K/Ar determinations yield estimates around 850 - 900 ka (Cox, 1969). Watkins (1968) recognized a younger, 830 ka, excursion. The excursion probably includes Nzone Zykh with ash giving a K/Ar age of 820 -t 250 ka in the upper Apsheron beds at Lake Zykh, Azerbaijan (Zubakov and Kochegura, 1973). According to new data recently reported by Mankinen and co-workers (1981), the K/Ar age for the Jaramillo event ranges from 900 to 970 ka, but the above authors also reported the presence of N-zone Taylor Valley with K/Ar dating at 840 ? 30 ka. Thus, in addition to the Jaramillo zone (nlb), the Zykh -Taylor Valley zone (nla) having an age range of 830-850 ka probably exists as well. (10) A double N-zone Kverno-Natanebi with age estimate at 1 .O- 1.1 Ma was recognized by Zubakov and Kochegura (1971, 1973) near the ChaudaIGuria boundary, Tsvermagala Mountain. Fleck and co-workers (1972) recorded an unnamed N-zone dated at 1.05 Ma from K/Ar analysis in lava sheet, Argentina. A 1.1 Ma fission-track age was determined on ash from N-zone in a lacustrine sequence of the Osaka Formation, Japan (Maenaka et a., 1977). Two excursions assigned to the Kvemo-Natanebi zone with age estimates at 1.O and 1.1 Ma were reported by Bucha (1976) from the oldest loess in Czechoslovakia. The most comprehensive data on the event were obtained by Mankinen and co-workers (1978, 1981) in California for the Cobb Mountain lava dome and for Clear Lake basalts. Here the duration of Nzone, named the Cobb Mountain, is estimated to be about 10 ka. The K/Ar analysis yield an age of 1.12 k 0.02 Ma.

Resume (1) Since the publication of the work by Shackleton and Opdyke (1973) time subdivision of the deep-sea Pleistocene sequences has totally given way to the oxygen-

' Two overlying R zones - Snake 11 and Snake I

- in the section, from lava flows 5 6 and 8, with K/Ar datings at 227 f 30 (average of three estimates) and 230 k 85 ka, probably correspond to the double Levantine - Tsokur event. ~

66

isotope scale. Micropaleontological, carbonate, oxygen-isotope and carbon-isotope cycles are in good agreement with each other, thus suggesting their common climatic nature. This allows their consideration as specific effect of global climatic events occuring in the Pleistocene with typical duration measuring tens of thousands of years. To unify climatostratigraphic terms the authors propose to introduce a term “orthoclimathem ” (OCT). Oxygen-isotope stages recorded in stratotype core V 28-238 and in parastratotype cores DSDP Hole 504, DSDP Hole 517, and M 13519 are used as stratotype of orthoclimathems. ( 2 ) Climato-micropaleontological cycles derived using transfer functions after Imbrie and co-workers (1973) are more sensitive to fluctuations in temperature of the surface water layer than isotope cycles. Therefore, the difference between them reported by some authors (Nilolaev, 1981; Blyum, 1983) is an accepted fact which needs further investigation. The above authors are likely to be right that micropaleontological cycles are in better agreement with the traditional notions of five coolings occuring during the Pleistocene. So far the notions have been substantiated insufficiently and inconsistent with conclusions drawn by Kukla (1977) and many other authors who found a striking similarity between isotope curves and soil and loess series in terrestrial sections. (3) To increase the validity of the climatic “ocean-continent’’ correlation it is a good practice to use an independent magnetochronological control and, particularly, data currently available on the geomagnetic polarity excursions.

SECTION I I

EVIDENCE FOR CLIMATIC CHANGES IN THE PLEISTOCENE REGIONAL REVIEW

As stated in Section I, the Pleistocene marine sequences are much more complete as compared with the continental evidence, hence the absence of more or less substantiated unified chronostratigraphic scale for the continents. Many local and regional schemes were constructed which are difficult to correlate. Therefore, we should consider original data currently available for certain regions to get a true succession of Pleistocene climatic events on the continents. In our review, not predetermining the question whether the isotope scale coincides with terrestrial data or not, we use it for comparison as a scheme tested in action for the whole of the world ocean. Below is a review of twelve regions on which the authors obtained more or less systematic information currently available on the climatostratigraphy of Pleistocene sequences and Pleistocene climate. Naturally, the succession of climatic events over the territory of the USSR is reviewed at greater length, than that of other countries.


Chapter 3 EFFECTS OF GLOBAL CLIMATIC EVENTS IN THE MEDITERRANEAN - CASPIAN SYSTEM

3.1. The Mediterranean as a new climatoparastratotype region Traditionally, the Mediterranean has always been the type area for chronostratigraphic division of the Pleistocene. However, if no stratotypes were designated and no reliable correlation with fossiliferous sections on the plains was provided for morphostratigraphic units of the northern Alps (Mindel, Riss, Wiirm), for the southern Alps correlation seems better substantiated. The Wiirm and Riss complexes of terminal tills which dam Lakes Como, Garda and others, are tied in to red soils (“ferreto”) and lacustrine-alluvial deposits drilled in the Leffe brown coal field, and are confirmed by pollen and spores of thermophilic associations and bone remains. An age of about 740 ka for the Giinz fluvioglacial gravels in the Rona basin was from K/Ar isotopic age measurements on overlying ash from the Agde volcano (Monjuvent et al., 1984). De Lamothe, Deperet, and their followers recognized glacio-eustatic raised beaches, related to the Alpine interglacials, in the coastal sections of the western Mediterranean (Zeuner, 1959, 1965). Erosional phases of the Apenninens were also related to the Alpine glacials (Blanc, 1957). Thus, the traditional Alpine terminology pertinent to the Mediterranean is the most valid in terms of stratigraphy but, unfortunately, this is a subject of controversy between different authors (Liimley, 1968; Bonifey, 1975; among others). In the last few years the traditional schemes have been revised. The marine morphostratigraphic division of Deperet and De Lamothe has been virtually supplanted by local stratigraphic nomenclature before there was any means of correlation even for the western Mediterranean. Table 3.1 presents a synthesis of Pleistocene climatostratigraphy for the Mediterranean based on data available from different sources, along with correlation of three types of divisions, namely raised beaches, bioclimatic zones by fossil mammals, and isotope stages from analysis of deep-sea cores. Up to 15 Pleistocene raised beaches with elevations varying from place to place are developed along the Mediterranean shore. Biostratigrahically, the raised beaches are divided into only two stages, namely, Sicilian and Tyrrhenian. The Sicilian is characterized by the occurrence of cold-water North-Atlantic emigrants Arctica islandica, Panopea norvegica, Hyalinea baltica, etc. The Tyrrhenian yields tropical Senegalese species Strombus bubonius, Tapes senegalensis, Conus testudinarius, etc. Some workers recognize also the Milazzian as an intermediate stage with moderately thermophilic molluscs Patella ferruginea, Tapes rhomboides, etc. (Blanc, 1957). According to the new terminology accepted in Italy the stages are equivalent to the Portuensian, Tarquinian, and Strornbus (Ambrosetti et al., 1972, 1978). The chronology of the raised beaches was worked out in great detail by Butzer for Mallorca (Bowen, 1979).

70 Table 3.1. Climatostratigraphic units of the Mediterranean Pleistocene (KM thermomer, age lo3 yr) Deep-sea record

Mammalian

~

kryomer, TM

Bioclimatozones

Marine terraces and erosional phases (er ph.) Pontinian er.ph. 14C 58000

1

3rd (Epityrrhen)

{ 1

2

Th/U 125

+

?

Loisia - Regourdou. Dicrostonyx

So00

.G

Y ~3m , Th/U 110 k 5 F.I. 90 + 18

Combe Grenal

1 2 a” . mc c

10, F.t. 127 f 13

~

0 5

4

Sanreney “B” KM Microtus oeconomus Santeney “A” TM Eliomvs - Clerhrionomvs Fontechevade K M L . Iaxurus - C. cncelus

Gr. Prince 1st

(Eutyrrhen)

I

F.I. 177

* 30

Grimaldi TM Eliomys

x , , 2-5 m Th/U 210.

La Fage KM Ostian erosional

Dicrostonyx - Lemmus

Orgnak “3 C” TM Orgnak “3 H” - Arago K M Dicroslonyx Ior(/uatus Orgnak “ 3 I” ( = Perrieres)? TM Hystrix major

F.t. 280 & 30

Riano? w 2 -

~

I ? - 23 m

Nomentanan (“Riss I ” ) er.ph.? K/Ar 430 Tarquinian, 40 rn = wI, 3 0 - 3 5 m Torre in Pietra. K/Ar 438 ? 40

1 =1

Pario‘i

“Mindel 11” er.ph. I = Nomentanan?) v2 - I 5 m Millazzian

Ranuccio

> 365

-

< 48

Flaminian erosional phase ( = Mindel 1.’)

Escale “G” Arvirola - Macacus

I

M c e 3 - Escale “F-C” KM Dicrostonyx Piiymys gregaloides

a

/ / / /

/

Vallonnet St Prest? Valerots 111)

I=

All pl.pitymoides, Hysrrix, Apodemus

invstocinus

__~ Valerols I I KM. A / / . nuhens Sicilian I - Calabrian ul.9h-104m.ESR12CQ~

1

Dicrosronvx onitquirorrs

Member\ “D-E”

1

Mas Ramboult - Valerots I lmola - Sinzelles. KlAr 13W

-

71

In Table 3.1 raised beaches are placed according to K/Ar datings and fissiontrack measurements for Italy (Ambrosetti et al., 1972, 1978) and U-series measurements for Mallorca (Bowen, 1978). According to new data (Ruggieri et al., 1984), the age of the base of the Sicilian in its type area at Cava Puleo of Ficarazzi, near Palermo, relates to the disappearance of Helicosphaera sell; and, hence, is about 1,150 ka. This is a traditional lower boundary of the Pleistocene and as such includes (after Butzer) no less than seven littoral sedimentary cycles A - G, each being subdivided into terrace levels, Z - U, and eolinites. The Sicilian comprises the three cycles G , F, E, and the Tyrrhenian consists of the four cycles, D, C, B, A. Italy and France represent an area where mammals both large (Bout, 1970; Azzaroli, 1983) and small are well studied. The study of rodents yields especially valuable paleoclimatic information (Chaline, 1977, 1978; Monjuvent et al., 1984). The recognition of rodentian faunal assemblages is based on the ecologo-climatic principle with consideration not only for the evolution of rodentia, but also for changes in the species ranges due to climatic variations. Moreover, in Chaline’s schemes biozones are named “climatozones”. For the last 1 Ma Chaline recognized 18 “climatozones” with alterations of kryomeric and thermomeric rodentian assemblages. The kryomeric assemblages are characterized by the presence of emigrants from Asia with tundra species including O b (Lemmus) and bog (Dicrostonyx) lemmings; the thermomeric assemblages give evidence for the abundance of Mediterranean forest species typical of maquis ( = gariques, sylvan), such as porcupine Hystrix, southern forest bank vole Clefhrionomys, southern arvenicolous hamster Allocricetus, garden dormouse Eliomys, etc. Invasions of periglacial Asian glacial rodents reached the Bay of Biscay and Provence and then retreated to give way to inhabitants of southern broad-leaved rather moist forests which advanced far north outside the Mediterranean. The eastern basin is the type locality for the climatostratigraphic subdivision of the Pleistocene Mediterranean deep-sea sequences. About 40 cores recovered from the area have been studied in detail using different methods; A (Albatros) - 189 is a classic core described by Emiliani in 1955 and Parker in 1958 (Cita et al., 1977); core RC 9 - 181 is the stratotype for a number of excursions (Ryan, 1972), etc. Sediments are subdivided on the basis of oxygen-isotope stages. Sixteen isotope stages were established by Verngaud and Grazzini in deep piston core KS - 09 (Cita et al., 1977). TypicaI of the deep-sea cores taken in the area are regional sapropel and volcanic ash horizons. Twelve sapropel and 23 ash layers have been established for the last 0.5 Ma (Cita et al., 1981); along with isotopic and micropaleontological data. This enables us t o get a high-resolution time subdivision for the Middle and Upper Pleistocene Mediterranean marine sequence (Parisi and Cita, 1982; Thunell and Williams, 1983, 1984; Muerdter and Kennett, 1984; among others). Core material which yields important paleoclimatic record is discussed in detail in section 3.4. It should be noted here that most of the sapropel layers are associated with thermomeric isotope stages, while ash layers are generally related to kryomeric stages (Fig. 3.1). Most of the sapropel layers and the thickest of them are related to isotope stages 5 and 7 and, hence, can be correlated with the Strombus within the raised beach succession and the interglacials of the Macedonian section (Wijmstra, 1978).

72

In the deep-sea cores the base of this warmest interval is marked by the appearance of Emilianiu huxleyi at about 240-220 ka (Raffi and Rio, 1979; Parisi and Cita, 1982). Comparison of the three cores suggests the presence of eight distinct climatic cycles for the Pleistocene Mediterranean sequence. However, in some cores the kryomeric and thermomeric parts of a cycle can be subdivided into a greater detail. The following is a brief discussion of the climatic features of the cycles. The first cycle of the traditional Mediterranean Pleistocene (G after Butzer) is confined t o the Cassian regression which is a probable equivalent of the Gunz. It is characterised by Epivillafranchian fauna Valerots with abundant arcto-steppe elements, including Dicrostonyx antiquitatis, in the north (Chaline, 1977, 1978). The thermomeric part of the first cycle corresponds to the first, warmest-water phase of the Sicilian transgression which gave rise to an abrasion platform at the

rn 0

8"O 70. PDB

Climatic and nannofossils in ka Emi Lcania

hux Leyi

I

Acme 61

2

3

Emi 1i a m a

huxleyi

4 22 5

t

Cephyrocapsi oceani c u

1 .- - -

55, TL 4 3 - 51 Aurignacian - Gravertian

.B P

j

North-Western Europe

Poland

I -1

9

?

2

II

Karma”? KM. ’*C > 46.8 Brerup - Odderade- Saint Germain TM TL 83-88. I4C > 69.5

- 2

1

EB 9 : 2 ,L?w

G Z

-

^

Y

Th/U84-100

Eemian-Tychnowy TM, TL l l O - l 2 5 Levallois - Moustierian - Preszeletian?

,iAi

Warthe - Lausitz KM Th/U 115-146

Warthe KM. TL 142- 179

- 4 . .

ML*

&.g: -

~

0

lliord - R u p n ? - Karlicher TM

y 5 !!

Reburg - Lamstedt - Flaming? KM

Podlesie TM Saale- Amersfoort

-

Hameln KM

c

Domnitz- Wacken TM

Zbojno TM

Holstein - Hoxne

I

10/12

z

3: U;

11/13

Barkowice-Mokre TM TL 378 - 389, FCI/P 320 - 420

1 11

Anglian KM. K/Ar 4W M . rrogontherir - prrrnrgenius Mundsley TM. Arv. cantiana Ferdynandow TM. TL 483.7 West Ranton TM

I Mogielanka KM. TL 532 - 544 Pilczyca TM, TL 558 San KM

I I 4 I l9

Kozi-Grzbiet TM, F W P 550- 700. TL 420 - 440

[ i e l m e .(“Glacial

Nida KM, TL 661-732

A”) KM

Artern - Osterholz TM. Last Eucornrnru

Przacnysz TM. TL 615-785

Elbe- Dorst KM

Unstrut

Narew KM

Lerrdani TM Linge KM

25 26-28

I 1 S‘elestynow TM?

1

( = &ten’, TL

Oiwock cooling?

I

923

*

1061

B o r n t a l ~Bavel TM Menapian I= Pinnow?) - “Gunz” KM

^

+

$5

Melisey I KM -Stump Cross

I

c

ij :

m

I

I

134

Wiegank (1982), Woillard and Mook (1982), Menke and Tynni (1983), Stremme (1983), Sutcliffe (1985) and nine Reports on the International Geological Correlation Programme Quaternary Glaciation in Northern Hemisphere (1972 - 1983) for the Middle and Late Pleistocene. Although the Soviet and foreign geologists have only limited information on Poland, it merits the greatest attention. The Polish researchers were among the pioneers in the development of the principles of climatostratigraphy. Rozycki (1964, 1969, 1982) and his proponents have greatly contributed to this field. An integrated approach, including thermoluminescence and fluorine - chlorine - apatite (FCl/P) datings, has resulted in detailed subdivision of glacial, loessal and cave deposits in Poland. The materials of Glazek and co-workers (1976, 1980), Makowska (1982), Wysoczanski-Minkowicz (1982), Lindner (1982, 1984), Lindner and co-workers (1982, 1983), Madeyska (1982), Mojski (1982, 1985), Lamparski (1983), among others are also listed in Table 5.2. Biostratigraphic, radiometric and magnetostratigraphic approaches were used in correlation of both columns of Table 5.2. Taken together, they gave a clear record of repeated climatic changes over central Europe between 55"N and 50"N. Except for stages and phases of the Last Glaciation, about 18 climatomers have been recognized for the last 1 Ma. For the purpose of our discussion we propose 16 major stages of climatic history, common for both north-western Europe and Poland (Figs. 5.2, 5.3). Ancient glacial time, named the Narew and Unstrut in Poland and in the GDR, respectively (Woldstedt, 1954; Lindner, 1984) and correlative with the Giinz of the Alps, is a complicated stage. In Poland, north and south of Warsaw, the Wyskkow and Przasnysz sections contain two or three tills separated by sands (Fig. 5.2); the tills overlie the Pleistocene Kozienice formation (Mojski, 1982, 1985; Lindner, 1984). In the Netherlands and FRG (Fig. 5.3) the ancient glacial stage consists of three coolings, namely Menapian ( = Pinnau), Linge ( = Elmshorn?), and Dorst ( = Elbe) separated by the Bavel and Leerdam ( = Pinneberg?) thermomers. During these thermomers, north-western Europe was covered by coniferous and broadleaved forests with Carya, Pterocarya, Tsuga and Eucommia (guttapercha tree); the latter now grows in southern China. Summer temperatures were then 2 - 4" higher than at present (Menke and Behre, 1973; Zagwijn and Doppert, 1978; Zagwijn, 1985). The second complicated stage, composed of two (or three?) thermomers corresponds t o the Ville pedocomplex, which lies on Rhine Main terrace I11 (Brunnacker, 1979), and to Cromer-I - I1 - 111 in the Netherlands (Zagwijn et a]., 1971). The first and strongest warming is known as the Podlasian (Rozycki, 1969) - Przasnysz (Lindner, 1984) in Poland, the Atern in the GDR (Erd, 1970), and the Osterholz in Denmark (Zagwijn, 1975). The warming was characterized by the development of oak and alder forest with rare Eucommia surviving. Two subsequent warmings - Mahlis I and 11 in the GDR - were weaker, but the findings of Hystrix in the Stranska-Skala beds (HoraEek, 1981) suggest a warmer climate than at present. These warmings are separated from the Atern by the Helme cooling, whose termination coincides with the Brunhes - Matuyama reversal (Zagwijn et al., 1971; Wiegank, 1982).

Y-ArAGE

PALAEOMAGNETIC POLA-

IN M.Y. R I T Y EPOCH EVENT

0-

SUBSERIESSERIES S T A G E p,

ICE-SHEETS NePOLANO -+S

AGE 1NM.Y.

AGE IN TH.Y

North-Polish

iOLOCENE

Middle-Polish

Mazovlan 0.5-

10.4

: : : I --

South-Polish

1.0.

-;:"0

\

20

40

GRUDZIADZ NTERSTADIAL

a

z a 1.5.

Krasnystaw -?-

Kozienice

1.83

-

3

c

213-

5

PRE-

GRUDZIADZ STADIAL

t

171

2 .o.

I

i

Fig. 5.2. Climatostratigraphy of the Pleistocene in Poland (after Mojski, 1985, fig. 4)

t

ProIUnslTL

I

I W i E R Z l N l E C I1 L

1

do

KASZUBY STADIAL (003NOwi.LI

EEM

I

,

136

New data reported by Polish scientists suggest that the South Polish - advance, ice extended along the Vistula river valley as far as the Nida river (50'30' N). TL datings of the Nida kryomer falling in the range of 660-732 ka and a reversed - normal polarity transition allow correlation with the Lika ( = Pokrov) stage and isotopic stage 20 (Lindner, 1984). A maximal ice sheet advance took place during the San glaciation, when the glacial edge abutted against the foot of the Car( = Cracow) - Elster I glaciation was three-fold. During the first - Nida

M E A N TEMP

YrGNEllC P4LEO-

IN

SCALE

106y.

BP

0 LATE PLEISTOCENE

JULV

10

20% J

5

Eemion

a

3 MIDDLE PLEISTOCENE

0.7 09

Holsteinion

a

>

"Cromerion

J -

Cr I

3

L E E R D A H INTERGLICIAL

3

I A V E L INTERGLIC(AL

Menapian

Eburonian

EARLY PLEISTOCENE 18 Tiglion

Proetiglion

25

LATE PLIOCENE

Reovenon

Fig. 5.3. Paleoclimatic curve of the Netherlands Pleistocene (after Zagwijn and Doppert, 1978).

137

pathians, at 49"30'N. The San and Nida advances in Poland are separated by the lengthy, but relatively cool (Malopolanian) interglacial (Lindner, 1982). Its lectostratotype is a section at Cave Kozi Grzbiet in the Holy Cross Mountains, 50"N, near the town of Kielce. A complex study of the section (Glazek et al., 1976) revealed three optima, represented by travertines, in the interglacial. The travertines contain thermophilic molluscan fauna (Helicigona banatica, and the like), amphibians, and reptiles. The mammalian fauna includes both forest (Ursus deningeri, Castor fiber, Pliomys lenki) and tundra-steppe (Lemmus lemmus, Dicrostonyx simplicior)

GDR after

A.G.Cepek (1967). A.G. Cepek ~ a l(1975; . J. Glazek sal(l98O

80

6

I

TREENE-WARMZEIT

.8

6

10

1 12 1 ELSTER KkLTZElT

'4 1

@

Fig. 5.4. Main climatostratigraphic units of the younger Middle Pleistocene in Poland and their correlation to GDR and USSR (after Lindner and Grzybowski, 1982). 1 - till, 2 - loess, 3 - localities of organogenic sediments, 4,5 , 6, 7, 8 - chronostratigraphic extent: of the Zbojno section (4),Wachock ( 5 ) , Swwiety Piotr (6), Karsy (7) and Kozi Grzbiet (8). 9 - samples with FCI/P and TL datings.

138

animals. FCWP datings on bones, ranging from 550 to 700 ka, allow correlation with isotopic stage 19. The San glaciation in south Poland consists of three stages separated by two Kielce and Pilczyca - warmings. Their interglacial rank has not been ascertained as yet. The TL method gives an age of 558 ka for the Pilczyca thermomer and an age range of 544 - 532 ka for the third - Mogielanka ( = Koch) - stage of the glaciation (Mojski, 1982, 1985; Lindner, 1984). Lindner proposed to correlate it with isotopic stage 14. The next time interval, equivalent to isotopic stages 15 to 11, is traditionally recognized in Poland as the “Great Mazovian interglacial” (Rozycki, 1969). Its equivalents in north-western Europe are the Voigstedt - Cromer s.str. and Holstein - Hoxne interglacials (Mitchell et al., 1973), separated by Elster I1 ( = Anglian) glaciation. In the Netherlands, tuffs from the Laaher volcanoes of the Urk formation have given a K/Ar age of 400 ka (Zagwijn et al., 1971). Recently the Mazovian interglacial has been divided by the Polish workers into three thermomers, namely Ferdynandow, Barkowice-Mokre, and Zbojno (Mojski, 1982, 1985; Lindner, 1984). Earlier Erd (1970) distinguished the separate Domnitz thermomer in the Holstein interglacial (Fig. 5.4). At present, sufficient data are available to consider the Ferdynandow, Voigstedt, and Frimmersdorf thermomers as synchronous with the Belovezha in the USSR and the Cromer Forest beds of West Ranton in England. All of them contain forest fauna with Palaeoloxodon antiquus, Alces latijrons, Macaca, Hippopotamus, Mirnomys savini, and the like, and yield pollen spectra typical of the coniferous and broad-leaved forest having oak, hornbeam, fir and such exotic species as Tsuga, Zelcova, Celtis and the like. Voznyachuk (1965) was the first to arrive at this conclusion. Apparently it was a two-fold interglacial. The first optimum in England (West Ranton TM) is marked by elm and oak forests and Biharian fauna with Mimomis savini and Hippopotamus; the second optimum (Mundsley - Westburry TM) is represented by hornbeam forests and by the presence of Arvicola and Macaca (West, 1968; Sutcliffe, 1985). During both optima climate was warmer and more humid than at present. The Elster I1 - Wilga - Anglian glaciation marks the disappearance of the Biharian fauna and the first appearance of a periglacial assemblage containing the early mammoth fauna with Mammuthus trogontherii, Ovibos cf. moschatus, Coelodonta antiquitatis, Rangifer sp., Dicrostonyx sp. Their remains have long been known from subtill pebblestones, including those from the type sections at Sussenborn, GDR, and in West Ranton. In addition to K/Ar ages at 400 ka and TL ages at 456 ka, this provides a reliable correlation of this glaciation with isotopic stage 12 or 14. Ice sheets were less extensive during Elster I1 ( = Wilga - Oka) time, as compared with those of Elster I and Saale time. Nevertheless, the first appearance of periglacial mammoth fauna is associated with this boundary. This can be easily explained if we assume that at time the climate in Europe was drier and cooler. The permafrost and kryosteppe zone was wider and the cooling lasted longer. The Holstein ( = Hoxne) interglacial can be reliably correlated throughout Europe (except England) because of the dispersal at that time of the Steinheimian

I39

forest fauna with Palaeoloxodon antiquus, Bos primigenius, Megaceros gigantheus, Macaca sp., and such evolutionally new rodent species as Arvicola cantiana, Pliomys episcopalic, Sorex savini, Cricetus cricetus, and the like (HoraEek, 1981). The Holstein interglacial is also characterized by a coniferous and broad-leaved forest (Fig. 5.5) with such rare exotic species as Tilia platyphyllos, Azolla filiculoides, Pterocarya, Celtis; very abundant dark coniferous species (Abies and Picea), as well as yew (Taxus baccata), holly (Ilex aquifolium), beech (Fagus Lu

B

ESTIMATED

MEAN TEMPERATURE

I N JULY

VEGETATION AND PaLEOCLIMATIC INOICaTDRS

E

1

B

O

U

L

D

E

R

C

L

A

Y

I

PROBABLY VERV OPEN LbNDSCbPE

OR POL4IiDESERT. PERYAFROST

OUERCU5 DlYlNlSMES rrSQ . ICarpinvr and AblDI

FORE-5

WITH TAXUS

OF

Ho 2

ALNUS.PINUS. RUERCUS PICEA,LITTLE ULHUS

a

auEi?cus

PINE'FOREST,LIlTLE ALN'IS.SOYE B E N L A

SU0ARCTlC P-KLANDSCePE IPINUS, JUNIPERUS, HERBS1

- - - - - _ _ _ - _ PE

LACUSTRO-GLbClAL CLAY l"FUIKLEI"1 REWORKED TERlIbRY POLLEN

Fig. 5 . 5 . Vegetational succession of the Upper Elsterian, Holsteinian and Lower Saalian of the Northern Netherlands (left column) and estimated changes of mean temperatures from Elsterian 10 Saalian times (right column) (after Zagwijn, 1973, figs. 1 1 - 12).

140

silvatica), i.e. the species which cannot survive cold winters (below - lo), but easily withstand cool summers (Erd, 1970). The areal distribution of vine embraced the British Isles, Denmark, and Poland. Frenzel (1967), who analyzed the specific composition of the Holstein forest, believes that during a climatic optimum in northwestern Europe the mean January and annual temperatures must have been, respectively, 1-3°C and 1-2°C higher, while in eastern Europe winter and annual temperatures may have been, respectively, 5 - 10°C and 3 - 6°C higher than at present. Precipitation must have been exceeded the present values by 50- 100 mm. According to Frenzel’s calculations and an independent estimate of Zagwijn (1957, 1973), July temperatures in western Europe were probably 2°C higher than at present. Hence, the Holstein is a pronounced interglacial with oceanic climate. However, offshore ocean waters and those of the North and Baltic Seas were no warmer than at present, as evidenced by marine molluscan fauna. The ranges of TL, FCI /P, and U-series ages for the Holstein are respectively 378 - 389 ka, 320 - 440 ka (Lindner, 1984), and over 350 ka (Stremme, 1983), but the electron spin resonance method (ESR) yielded ages ranging only from 200 to 240 ka, as reported at the XXVII IGC in Moscow (Abstracts . . ., 1984). The author thinks that all the above estimates are minimal. The U-series value exceeding 350 ka obtained by Mangini on shells of Littorina littorea (Stremme, 1983) seems to be most reliable. The Fuhne kryomer and its Polish equivalent - Liwiec - with a TL age of 352-368 f 44 ka (Lindner and Grzybowski, 1982) is considered now as an independent “minor glaciation”. An ice tongue reached then the latitudinal Bug rivercourse at 52”N. Some authors (Rozycki, 1969; Zagwijn, 1973) regard the Domnitz ( = Wacken - Zbojno) thermochron as a first interstadial of the Saal- Middle Polish glaciation, whilst the others (Erd, 1970; Cepek and Erd, 1982; Lindner, 1984) consider it as a “minor interglacial”. Palinological data on the Pritzwalk type section, GDR, suggest a close compositional similarity between the then and present forest (Erd, 1970). This is also valid for the Zbojno section in Poland (Lindner, 1984). Consequently, the summer temperatures of the Domnitz interglacial may have been 1” lower than those typical of the Holocene climatic optimum. The Saale ( = Walston - Odra) glaciation was a multistade one. Liittig (1968) recognized seven advances, but Rozycki (1969) and Lindner and Grzybowski (1982) established no less than five stades. Maximal advances (stades) were the Hameln and the Kamienna in the FRG and Poland, respectively; they may well have taken place at different times (Table 5.1) and could be separated by the Hoogeveen - Podlasie warming, when birch forests with oak and hazel grew in the Netherlands and Poland, and summer temperatures were only 3°C lower than at present (Zagwijn, 1973). During the maximal stade the Saale ice sheet locally straddled the Elster - San glaciation boundary and reached the Sudets piedmont in the territory of the GDR and Czechoslovakia. The time interval between the Saale and Warthe glacier advances and the age of the Eem marine transgression are the most debatable problems in the Pleistocene stratigraphy of Europe. Zeuner (1959) believed that the Eem interstade preceded the Warthe stade. However, Woldstedt (1958) showed that in the lower Elbe River area the Eem deposits intrude the marginal belt of Warthe tills. In the post-war years,

141

Picard, Cepek and other scientists recognized two, in their opinion, intra-Saale warmings, namely, Rugen and Treene (Cepek and Erd, 1982). Based on the new data Kukla (1977) suggested the presence of three Eem sequences: above the Warthe (Eem = Scaerumhede), below the Warthe, and within the Saale (Eem = Eem). He correlated them with isotopic stages 5, 7, and 9, respectively. Bowen (1978) and Sutcliffe (1985) came t o the same conclusion about the Ipswich. Three interstadials (Vejby I and 11, and Oksbol) in the Saale complex were established in Denmark (Lundqvist, 1982). Recent studies of the loess successions in Normandy (Lautridou, in Quaternary Glaciations. . ., 1982), in the Rhine river area (Brunnacker, 1979; Urban, 1983), and in Schleswig-Holstein (Stremme, 1983) also suggest the presence of one or two interglacial soils between the Eem and Holstein soils. In Poland the Odra and Warthe glacial advances were undoubtedly separated by an interglacial. It is evidenced by soils on the Lublin plateau with TL ages of 221 i 27 ka, lacustrine deposits with TL ages of 245 and 264 ka, and, primarily, by intertill alluvial deposits in the Vistula valley, near the village of Grabowka; here in the pollen diagram of a coniferous and broad-leaved forest oak pollen accounts for 20%. In the nearby Frombork section TL ages of these (?) sands fall in the range of 240 - 260 ka (Lindner, 1984). Classical sections at the Middle Palaeolithic Taubach and Ehringsdorf sites in the vicinity of Weimar provide evidence for two “Riss - Wurm interglacials”. Between alluvium and loess (“Pariser”) the section of each site contains two or three travertine horizons with numerous leaf imprints of broad-leaved trees, remains of animals and molluscs which dwell in the more southerly regions, as well as remains of Heidelberg man and Middle Palaeolithic (Mousterian - Tayacian) flint implements. Since Soergel and Behm-Blanke, who first studied the travertines, their assignment to the “Riss-Wurm” and correlation with the Eem transgression used to be thought unquestionable (Woldstedt, 1958; Zeuner, 1959; Markov et al., 1965). However, a new paleontological investigation showed that the Taubach and Ehringsdorf travertines differ in age. Thus, tooth histology, supported by U-series dating, suggests that Arvicola belongs to two different types. The ages of the Taubach travertine lie in the range 93 t 16 ka to 116 k 23 ka, whilst the lower Ehringsdorf travertines dated by various methods have age ranges between 205 t- 90 ka and 220-262 ka (Heinrich, 1982; Jager and Heinrich, 1982). This allows the Ehringsdorf - Grabovka thermochron be considered as an interglacial intermediate between the Eem s.str. ( = Bobbitshole) and the Domnitz. After Kukla (1977), Bowen (1978) and Sutcliffe (1985) we tend to believe that it may have been the Early Eem ( = Ipswich) interglacial. Moreover, the data currently available suggest a more complex correlation, i.e. the Treene, Eem-Eem, and Ipswich-Ilford can be correlated with isotopic substage 7c, and the Rugen - Ipswich-Brandon with a Useries age of about 174 ka with substage 7a. The stratigraphic position of the Hoxne beds in England, Bilzingsleben beds in the GDR, and Kerlich beds in the FRG remains debatable and uncertain. On the one hand, these beds contain fauna of the Holstein type (Arvicola cantiana, Paludina, Corbicula fluminahs, and the like) and flora (Potamogebon filiculoides, Taxus baccata, and the like); on the other hand, they comprise traces of man assigned t o a stage intermediate between Homo erectus and H. sapiens, and traces of stone

142

implements which resembles the Middle Palaeolithic (Clactonian - Mousterian Tayacian) culture, and yield also Eemian guide fossils of molluscs. U-series ages at 150 ka for the Karlich (Urban, 1983), 228 f 17 ka for the Bilzingsleben, 245 k 25 ka for the Hoxne, and 272 ka for the Swanscombe (Glazek et al., 1980) suggest a post-Holstein age. They may well correspond wither to isotopic substage 7c, or stages 9 t o 11; U-series datings of the Holstein exceeding 350 ka (Stremme, 1983) may be underestimated. The Ehringsdorf - Grabowka interglacial, apart from the above-mentioned questionable sections, is characterized by a forest fauna with Palaeoloxodon antiquus, but such African elements as hippopotamus and hyena are absent, while Dicerothinus kirkbergensis, Equus and Arvicola cantiana-terrestris become very important (Bowen, 1978; Jager and Heinrich, 1982; Sutcliffe, 1985). Forests contained various broad-leaved and coniferous species, but relic forms, if any, are sporadic. This suggests that during isotopic stage 7 climatic conditions were similar to those at present. However, the occurrence of Thuja occidentalis and Rhododendron sp. in the Ehringsdorf points t o slightly milder winters. The Warthe glacial advance has left a well-defined belt of marginal formations in Europe. Different authors at various times assigned the belt either to the last glaciation (Zeuner, 1959) or t o the Saal (Woldstedt, 1954). New data indicate that the Warthe glacial advance may have been an independent glaciation. TL measurements on drift and loess have given respective age ranges of 147 - 156 ka and 142.5 f 3.6- 179.7 k 22 ka for the glaciation in Poland (Lindner, 1984). Equivalents of the Warthe kryochron in the Ehringsdorf section are upper travertines, which include, according to Kahlke, Mammuthus prirnigenius and Coelodonta antiquitatis. Their U-series age is about 146 f 30 ka (Jager and Heinrich, 1982). The Warthe glaciation was less extensive as compared with the Saal s.str. The Eem s.str. - Ipswich s s t r . interglacial is associated with a warm-water transgression when on coasts of the Baltic Sea and north-western Europe the appeared Lusitanian molluscs with Tapes aurea var. eemensis, which now live off Portugal. Taking into account the uncertain extent of the Eemian beds, the Fjosanger section at Bergen Fjord, Norway, which has been recently studied in detail by Norwegian scientists (Mangerud et al., 1981), can be regarded as a reliable parastratotype. Pollen zones, typical of the Eern optima, were established there in marine sands with shells dated at 130- 140 ka by the amino-acid method. A similar section has been studied in the lower Vistula area, where the Eemian fauna occurs in two horizons. The upper, Tychnowy, horizon contains 36 molluscan species, including Paphia aurea senescens, Eulimella nitidissina, etc. ; its pollen diagram corresponds to an Eem Interglacial optimum. The lower, Sztum, horizon with poor fauna (Cardium, Nassa, Mactra, and the like), separated from the upper horizon by barren silts (Makowska, 1982), probably corresponds t o the Ilford. We recall that the subdivision of the Eem transgression into the lower Eem with Turritella and the upper Eem with Abra alba in successions of the Netherlands and Schleswig was suggested by Wolfe as early as the 1930s and later by Heide (1965). The continental facies stratotypes of the Late Eemian (Eem s.str.) interglacial, i.e. of isotopic substage 5e, are the above-mentioned Taubach travertines with Bums sempervirens and Arvicola cantiana-terrestris, dated by U-series around 93 - 1 16 ka,

I43

and travertines comprising Hippopotamus amphibius and Dicerorhinus hemitoensis in the Yorkshire caves with eight U-series datings ranging from 114 i 5 to 135 i 8 ka (Gaskoyne et al., 1981). Since the studies of Jessen and Milthers (1928), six pollen zones have been traditionally established in the Eem, although their diagram was composite. Based on the Grande-Pile bog in the Vosges Woillard has recently constructed a unique complete diagram embracing a 140 ka long time interval (Woillard and Mook, 1982). In the Eemian climatic optimum (phase "f" of Jessen and Milthers, 1928), in central Europe grew a broad-leaved forest with oak and yew, which later gave way to a forest mainly with hornbeam and fir. The reconstructions of climatic conditions for that time were made by Frenzel (1967), and later by Grichuk (Gerasimov and Velichko, 1981), and Velichko and co-workers (1982). Frenzel and Grichuk received independent A T, for north-western Europe 1 - 3°C and 1 - 2"C, and for Poland 3°C and 4"C, respectively. A T, obtained by Frenzel are 1 - 2°C and 3 - 4"C, while those of Grichuk reach 6 - 7OC; both obtained the same annual A T estimated at 2 - 3°C. According to Frenzel, precipitation exceeding the present level by no more than 50 mm (for Poland), but Grichuk presented a value of 200 mm. However, the temperature of ocean surface waters off Europe differed even more from that at present. In travertines of the Eemian 8 m-high terrace with U-series age at 123 +_ 24 ka on Jersey Island, Keen and co-workers (1981) found a mollusc Astraliurn rugosum, which now does not occur north of La Rochelle. This yields annual A T for offshore waters around 3 - 4°C. North-western Europe is a type area for minute chronostratigraphic subdivision of the last glaciation. By tradition, the Vistula kryomer is divided into three parts (Table 5.2). The type sections of the Early Vistula subkryorner are the Brorup section in Jutland and the Amersfoort in the Netherlands, where early Vistula time has been subdivided into four stades, separated by the Amersfoort, Brorup, and Odderade interstadials (Van Hammen et al., 1967). More complete sections of the Grande-Pile bog in the Vosges and of the Keller Borehole in Schleswig-Holstein have recently been studied in detail. The revision of the Early Vistula pollen diagrams (Menke and Tynni, 1984) showed that the Rodebek and Amersfoort interstadials are fragments of the Brorup interstadial, and that only two interstadials, namely, the Brorup, synchronous with Saint-Germain I , and the Odderade, synchronous with Saint-Germain I1 occurred in Early Vistula time. The conclusions of Menke are consistent with new data on Poland (Rozycki, 1982). There are some complete loess sections with buried soils and Palaeolithic cultural layers there. A series of TL and 14C datings on the Zwierzyniec section, near Kracow (Madeyska, 1982) make it the most important section. In western Europe the Saint Romaine section, near le Havre, also dated by TL, has been designated as the type section for the Late Pleistocene (Wintle et al., 1984). New data allow subdivision of the Early Vistula kryomer into five parts, common for the entire area studied (Table 5.2). The first kryostage corresponds to the Kaszuby loess of Poland with TL datings ranging from 97.7 to 101.3 ka, and with Blake excursion (Mojski, 1982, 1985; Lindner, 1984), to the Melisey I cooling and isotopic substage 5d. It is believed that as early as that time the Scandinavian ice may have filled the Baltic Sea basin with its tongue entering into the Vistula river

I44

mouth up to Malbork (Mojski, 1982). A U-series value of 110- 84 ka was obtained for travertines containing remains of Gulo and Rangijer in Stamp Cross Cave, Yorkshire (Sutcliffe, 1986). During the first Vistula interstadial, known as the Brorup (Menke and Tynni, 1984) -Saint Germain I (Woillard and Mook, 1982) - Jozefow (Dylik, 1968)-Fan (Mangerud et al., 1981), a pleasant climate like our own still persisted in northern Europe, as suggested by the distribution of a broad-leaved forest in the area. Even Lapland was free of ice. The second interstadial, named the Odderade (Menke and Tynni, 1984) -Saint Germain I1 - Chelford - Fornes (Lundqvist, 1983) is well represented in the Zwierzyniec section by a double buried soil with TL ages falling in the range of 72.9 - 75.9 and 71.6-72.2 ka (Madeyska, 1982). The 14C method yields an age of 69.5 + 3.8 ka for the termination of the interstadial in the Saint Germain I1 section (Grn 1987). Since the interstadials are separated by a minor cooling, in many sections the Saint Germain thermomer can be regarded as an entity represented by the Jamtland in Sweden, Perapohjola in Lapland, and Saint Romaine in Normandy. From TL datings varying from 83 f 7 to 88 f 8 ka, the Saint Romaine thermochron is correlative with isotopic substages 5a-c (Wintle et al., 1984). The double Saint Germain thermomer (“B” according to Kukla, 1977) is marked by the presence of Levallois - Mousterian and, possibly, Preszeletian (Madeyska, 1982) flint implements in the buried soils and in loess. The third cooling reached a maximum in Early Vistula time. The Torun till, recorded in the lower Vistula river area, may be related to the cooling (Mojski, 1982). This cooling corresponds t o the middle Vistula loess ( = SartowiEe) with a Late Mousterian and Micoquian - Preszeletian time interval and TL datings falling in the range of 67.6-71.7 ka in Poland (Madeyska, 1982) and 75 6.5-80 f 7 ka in Normandy (Wintle et al., 1984). The above estimates unambiguously indicate that the Middle Vistula kryomer was synchronous with isotopic stage 4 (73 - 61 ka). This is at variance with the validity of final I4C datings for the Brorup - Odderade and a “short-term” time scale for the Wurm. The middle Vistula pleniglacial, corresponding to isotopic stage 3, may be only tentatively referred to as a thermomer. Really it includes three buried soils, but they are of the tundra - steppe or podzolic type (“B3” after Kukla, 1977) associated with frost-pattern soils, suggesting the presence of permafrost in northern Europe during Middle Vistula time (Hammen et al., 1967; Dylik, 1968, 1969; Washburn, 1980). The first and most intensive warming known as the Upton-Warren - Moershoofd in north-western Europe and the Konin - Gniew - Maliniec in Poland, had left podzolic soils with first Late Palaeolithic implements of the Szelet - Jermanovician culture. 14C measurements yield different ages ranging from 23 to 45 - 50 ka, while TL ages for the Zwierzyniec section are around 47.3 - 50 ka. In the Grande-Pile section this warming corresponds to zone 14 with 14C ages below 49.8 ka. Coope (1977), who studied coleopterans (beetles) from the interstadial Upton-Warren beds, England, revealed a discrepancy between the results of spore and pollen analyses and specific composition of insects. The former suggest the development of forest - tundra landscape, i.e. a cooler climate than at present. Nevertheless, the beetle fauna is represented by species which occur now in Spain and in the Caucasus

i 45

and, hence, points t o higher summer temperatures, as compared with the presentday temperatures in England. The beetle fauna is able to mark only short-term, on the order of hundreds of years, climatic warmings, which could not affect the vegetation. The second, minor warming is represented by the Hengelo tundra - gley soils with 14C ages at 39 k 2 ka in the Netherlands (van Hammen et al., 1967), Grudziadz soils with an age of 40.7 -t 2 ka in Poland, and zone 15 with 14C dating of 40 _t 6 ka in the Grande-Pile section. The third warming was a two- or three-fold thermomer with its early peak reflected in the Arcy interstadial soil, France, with I4C ages of 31.5 - 30 ka. The middle and late peaks are represented, respectively, by the Kesellt soil with 14C age around 29 - 27 ka, and Tursac soil with ages of 24 - 23 ka. This three-fold interstadial, known also as the Denekamp (= Gota-Ah - Sandnes and others) can be recognized in many loess sections from the findings of stone implements of the Aurignacian - Gravettian type (Renault-Miskovsky and LeroiGourhan, 1981). The second and third interstadials are separated by a cooling, when the mouth of the Vistula River was again invaded by a glacier tongue which left the Swiecie till. Based on 14C datings, Vistula ice, which left the Brandenburg (Leszno) belt of terminal tills on the continent and the Devensian till in the British Isles, acquired a maximum extent not simultaneously, but within a time interval of 23 - 19 ka, mainly between 20-22 ka BP. The northern coasts of the Baltic Sea freed of ice during the Raunis interstadial, about 13.5 ka BP. Thus, the maximum of the last glaciation lasted only for about 10 ka. At that time periglacial conditions set up in middle latitudes of Europe. Treeless tundra - steppes on permafrost, active loess accumulation, and intense kryolithogenesis, studied in detail by Polish (Dylik, 1969) and Dutch (Van Hammen et al., 1967) scientists permit reconstruction of a rather severe, extreme continental climate with lowering of average July temperatures to 3°C within the periglacial zone. Hence, A T, measures 15°C for Late Vistula time of the Netherlands.

5.3. West Siberia

The West Siberian lowland with its network of south - north rivers presents a unique opportunity for climatostratigraphic correlations of Pleistocene deposits in various latitudinal zones, that is from tundra to steppes. Facies changes within five assemblages can be established in this gigantic meridional profile extending over 2500 kilometers in the shore sections of the Yenisey, Ob and Irtysh, these are marine-shelf, glacial, lacustrine-alluvial loessic, and mountain-glacial deposits (relationships between glacial and marine sediments will be discussed in Section 5.6). The composition of West Siberian Pleistocene deposits was described by the author in his earlier works (Zubakov, 1972a; 1972b: 1974). The following discussion of climatic episodes (Table 5.3) is based on the results of the author’s research, supplemented b y new evidence (Kaplyanskaya and Tarnogradsky, 1984, for the West Siberian lowland: Borisov (1 984) for the Sayan - Altai Mountains. New results were published by Arkhipov, (1984), Arkhipov et al. (1977, 1980, 1982), Volkov et al.,

146

(1984), Volkova et al. (1984), Grechin (1975), Zazhigin, (1980), Zazhigin and Zykin (1984), Svitoch et al. (1978), Zudin et al. (1982). Paleoclimatic interpretation of the data is presented in Fig. 5.7. The West Siberian lowland is the largest area in the Northern Hemisphere where lacustrine-alluvial deposits are widespread. It can be compared to a pancake on a gigantic plate, 1800 km wide. Valleys of the modern rivers are not larger than 100 - 200 km though, having three or four terraces above the flood-plain. Their age Table 5.3. Climatostratigraphic units of the West Siberian Pleistocene

7Area of - lake and loess sedimentation 1 Allai Mountains

4 a

~- -

[

1 'g I C

-1j1 1 I 4C5d

5e

1

river

Glaciation area

Norilsk Sopkey stade 1st terrace ("Karginian") ~

--

14r 7 1 - 7IQ, ~

.~

N'yapan Lokhpodgort Zhigansk? srade. I4C 34 40 lgarka TM, I4C 40 > 50 Ermakovo ( = Khashgort?) stade

Mirnoe-Karymkary TM

glaciation, KTL 110 f- 17-240 KTLIMTL 246

'-I0

1

f

30

Samarovo - Bakhta glaciation KTL230-312

Till with KTL 413 f 52 (Nizyam? KTL 420-510?)

i 8 1 1\ u

R-polarity, Belavo beds. KTL 610 k 70, Mezosiphnsus sp - Allophoromys

26 - 30

Mokhovrkaya suite - upper part Taman' fauna - Prol. proeponnomcus

r

n

Teleirk T.M, KTL 630 i 27 Fluvioglacial conglometate, R-polarity

Beken suite? KTL 910-1200

I47

is established as the Late Pleistocene. The interfluves expose generically peculiar strata showing lateral facies changes of the following sediment assemblages: lacustrine, lacustrine-alluvial, talus-alluvial, eolian, solifluctional. The sequence is marked by the alternations of kryogenic-congelation deformations of the strata, including pseudomorphs after polygonal-veined ice with hydromorphous soil topped by peat bogs or the alluvium of small rivers. The sequence described is similar to the deposites of the Jana - Kolyma lowland. The author recognizes here a new formation the so-called congelifluction-sor assemblage (paragenesis)2 considering a complicated paragenesis of facies and genetic types dominated by shallow water lacustrine and fluvial deposits in the permafrost environment and under the condition of periodic jams due to ice. Volkov et al. (1978) claimed the interfluvial deposits of the southern West Siberia accumulated in vast and deep affluent Mansi “Lake- Sea”, which appeared every time North Siberia was glaciated. Thus, the main point of disagreement is an estimate of lacustrine sediments in the interfluvial succession. The congelifluction-sor assemblage also fills ancient buried river valleys of West Siberia, their pattern being almost the same as the network of modern rivers there; the portion of lacustrine sediments in the congelifluction-sor assemblage increases northwards. In the vicinity of the Ob, Irtysh and Yenisey valleys it is intruded by interglacial alluvial assemblages. The total thickness of sediments in old buried valleys of the Yenisey reaches 100 meters (Fig. 5.6), it is much larger in the Upper Ob.

120 -

80 -

I

t i g . 5.6. Morphostratigraphic scheme of the Pleistocene and Upper Pliocene deposits of the upper Yenisey River valley (after Zubakov, 1972a). 1 gravel and sand, 2 - loam, 3 clay. 4 - loess and congelifluction-sor assemblages, 5 bui-ied 60. T h / U IW Kresta Name RiverEklutna? KM

Taiga with Corylus R T L 176 t 20

Mechigmen - 0 s s o r a Kotaebue TM. Va - IV foram-zones, RTL 184-204, T h / U 175 (233?)

Lower Vechernino - Lanzhin T M Taiga with Abies. Corylus RTL I12 28

Poluden - Okhota KM R T L 470 f I20 Belichan T M Taiga wirh broad-leaved trees Protochnino - Avlekit KM R T L 580 + 150 Delyankir - Ketanda T M Taiga with Tsugu, Juglans, Pin. Omoricu R T L 647 2 78 KM?

Achchagyt ( = Krcrt-Yuryakhl Thl north taiga D. simphczor-iorquorus

~

*

Jurov KM. Forest -tundra. R T L 250 f 50 Kyurbelyakh-Urak T M h n . Omorrca - Corylus RTL 350 t 87-410-478

Oiyagos KM, tundro - steppe

~

~

8

.-

5

8 ' 5

3

.-

Allaikha-Utka-Maastakh Khroma KM Wcrosionyx cf. simplrcror.

Equus cf abrlr, Mummuihus pnrnrgenius

Tundra. tundro - steppe

Olyaion- Caribou-Hills? K M Regression Yanrakinot - Einahnuhto transgression 70? K/Ar 320

Adycha? T M (Middle Ulakhan-Sular). North taiga. Lake Jakutskoe hcds

+

KuchchuguiLower Lllakhan-Sular suiie

Mitogtno KM

Dtcr renidens. Proeuke3 luiifrons. Norrh raiga

Karagino tranygreasion? llnd foram-zone: Elphrdiellu rolfi

Y

Pinakul? - KhomutaMount Susitna KM?

/

Praeovihos lllrd zone- Akan T M

1st foram-zone -

ClerhrronomyP

Elphidrello quostorego

North taiga, peat

nc"sIY

Upper Ol'khaTusatuvayam beds

I3 5

R polarity

Chukochiya beds 1st - llnd Oler hiorone UP^' Tigil - Iran-Creek KM

A

ig

illlopharomys c/ pirocoenrcus, Lemmw c l . obenrrr. Praeovrhoi, Prrdrcrorront 8 compirollu c

In addition to the above, the column “Mountains” is based on the data of Voskresensky and co-workers (1984), Bespalyi and Davidovich (1984) for the intermontane troughs in the Indigirka - Kolyma interfluve, 62 - 66”N, and the material of Ananiev and co-workers (1985) on the middle mountains in the north-western Sea of Okhotsk area, 140- 150”E. The mountain system in north-eastern Asia is known to have been covered with valley glaciers and ice cap and piedmont glaciers (Kolosov, 1947; Zamoruev and Petrov, 1984). The intermontane troughs were only partly ice-filled and, hence, the facies and stratigraphic relationships between the glacial assemblage and alluvial deposits of ancient valleys can be traced within the troughs. In the well-known Berelyekh river valley, a left tributary to the Kolyma river, all in all 17 terraces above the floodplain were established; they were grouped into four terrace complexes with elevations of 5 - 25, 40- 50, 110- 115 and 125-220 m. The terrace surfaces are overlain by thick trains of rhythmically laminated talus, in the sections of which buried soils and peat bogs containing fossil stumps and cones are interlayered with patterned ground, consisting of ice wedge polygons. The higher the terraces, the thicker and more complex are the overlying talus, which, along with the alluvial deposits, form so-called “terrace talus”. The sections of the latter were subjected to comprehensive palinological studies. In the Sea of Okhotsk area Ananiev and co-workers (1985) recognized six till horizons with TL ages ranging from 17 - 24 to 580 ka. They correspond to isotopic stages 2, 4, 6, 8, 12 and 16 (Table 5.4). The same number of tills and corresponding periglacial complexes were established by Voskresensky and co-workers (1984) in the terrace-talus sections of the Indigirka - Kolyma interfluve. Effects of an earlier cooling, probably corresponding to isotopic stage 22, have been revealed in the basal beds of the Ust” Delyankir thermomer. During three earlier interglacials, (Fig. 5.8), in the time interval ranging from 647 to 410- 350 ka BP (RTL estimates), the middle mountains were covered by a dark coniferous taiga bearing resemblance to present vegetation in the lower Amur river area, namely the middle mountains give evidence of abundant fir, broad-leaved trees (oak, elm, lime, nut), exotic fir and pines, Picea sect. Omorica, Pinus sect. Srrobus, and the like (Voskresensky et al., 1984). The extinction of dark coniferous taiga in the mountains of the north-east USSR started with the accumulation of 125-220 m high alluvial deposits of reversed polarity, correlative with the Matuyama ~ r t h o m a g n e t h e min~ the Berelyekh river valley, and terminated with the maximal left-sided Berelyekh - Yurov ( = Elga?) glaciation, which occurred, according to RTL measurements on till, about 250 i 50 ka BP (Ananiev et al., 1985). During two subsequent thermochrons with RTL datings at 176 k 2 ka and 112 28 ka, a light coniferous taiga with scarce fir-trees and the most coldresistant broad-leaved varieties (Tilia, Corylus) grew in the mountains of the northeast. The distribution of a light forest in the middle mountains suggests a climate similar to our own or somewhat cooler during warmings with RTL ages falling in the ranges of 37-44 and 17-24 k a (Fig. 5.8). The marine sequence was studied by Petrov (1966), Svitoch and co-workers

Bespalyi and Davidovich (1984) assign it to the Gilbert orthomagnethem.

IS')

(1980), Nevretdinova and co-workers (Biske, 1982), lvanov (1983) on the Chukotka coast; Gudina et al., 1984), Karetskaya and co-workers (1984) on Ayon Island; Gladenkov (1978), Svitoch and co-workers (1978), Bespalyi and Davidovich ( 1 984), Petrov (Zamoruev and Petrov, 1984) and others on the Kamchatka coast. The table also includes data reported by American scientists for north-western Alaska (Karlstrom, 1964, 1968; Hopkins, 1973; Hopkins et al., 1974; Weber et al., 1981). The above data point to the presence on the Bering Sea coasts of no less than five marine terraces with heights varying from place to place. The composition of molluscan and foraminiferal faunas and diatoms enable division of the marine sequence into two biozones, namely ( i ) the Pliocene - Early Pleistocene biozone, containing forms which do not live now i n the Bering Sea (molluscs Sw$ftipeccren .sw$fti, Asrarte invocata; benthic foraminifera Eiphidieiia hunnui, E. niridu, E. quasioregonensis Gud., Cassidulina luricainerafa, and the like; diatoms Melosira aibicans, and the like); and (ii) the Pleistocene biozone with organisms which are still living in the Bering Sea. The first zone comprises the Olkhovskaya formation in eastern Kamchatka, the Tusatuvayam beds on Karaginsky Island, the lower PinakuI beds of Chukotka, the Anvilian beds in Alaska, and the first Enmakay foraminiferal zone in the section drilled by a key hole on Ayon Island. Intervals with normal and reversed polarity have been recognized in the zone. The presence in the Anvilian beds of dextral tests of planktonic foraminifera Neogioboquadrinu pachyderrna, typical of the Jaramillo event, allows a tentavive correlation o f all these deposits with isotopic stages 21 to 19 or 25 to 19. The Skull Creek till in the Seward Peninsula, which underlies the Anvilian beds (Hopkins, 1973), may corre-

Fig. 5.8. Estimated climate of intermontane depressions of north-eastern Asia through the Pleistocene and Late Pliocene (after Grichuk in Voskresensky et a]., 1984). Q l v -N, - age symbols after Grichuk, (a) mean annual temperature, (h) duration of non-frohi period in days, (c) summer (July) temperature, (d) %,inter (January) temperature, (e) mean annual precipitation, (f) orthoclimatheni (after the author).

160

spond, hence, either to isotopic stages 22 to 24, or to stage 26 (0.8 - 1.15 Ma). The existence of glacial conditions in the montane area of Kamchatka at that time is evidenced from interbeds of iceberg till in the upper Olkhovskaya beds of the Kamchatsky Nos Peninsula. The Ca/Mg method yields water temperature values in the range of 3.6- 5.8"C for the Anvilian and lower PinakuI transgressions (Svitoch et al., 1980). The second transgression may presumably be associated with the formation of (i) the upper Pinakul beds, separated from the lower Pinakul strata by an erosion surface; (ii) the Karaginsky beds (Ivanov, 1983); and (iii) the second foraminiferal zone in the Ayon section (Gudina et al., 1984). Evidence of the third - Yanrakinot transgression was recorded by Ivanov (1983) in a 35 - 50 m high terrace of Kresta Bay; in the Lakhtakh formation of Kamchatka; and in the Einahnuhto beds of Alaska, which rest on tuff with a K/Ar age of 320 -t 70 ka (Hopkins, 1973). The Arcto - Boreal type of fauna characteristic of the transgression indicates a higher water temperatures than at present. In the Chuckchee Sea the transgression led to the accumulation of the third foraminiferal zone with Miliolinella pyriphormis. In Chukotka the terrace is overlain by till of the Olyaion glaciation (Ivanov, 1983). The fourth transgression, known as Kresta (Petrov, 1966) and Mechigmen (Ivanov, 1983) in Chukotka, Kotzebue (Hopkins, 1973) in Alaska, Ossora in Kamchatka (Zamoruev and Petrov, 1984), left a marine terrace with a maximal height of 30 to 35 m; it gives evidence of a typical Arctic fauna with Portlandia arctica and Elphidiella arctica. TL ages are about 184 +- 22 ka and 220 ka for the beds (Svitoch et al., 1980). A234U/23sU ratio with R = 1.0 and R = 1.15 gives for the Kotzebue beds ages of 175 ka (Karlstrom, 1964) and 233 ka, respectively. These estimates allow correlation of the Mechigmen - Kotzebue transgression with isotopic stage 7. Fauna and pollen spectra (Karevskaya et al., 1984) suggest that the climate that occurred at the time in the Bering Sea area was similar to our own. Water temperature values of 12.5 - 12.8"C obtained by the oxygen-isotope method on Astarte borealis placenta molluscan shell fragments from the Ossora beds of Kamchatka and that of 14.9"C on Macoma middendorfi valve fragments (Ivanov, 1983) appear unrealistically high. Till overlying the terrace is associated with the Krest - Nome River Glaciation, correlative with the Illinoian of North America (Karlstrom, 1964, 1968; Hopkins, 1973) and the Yenisey glaciation of Siberia (Zubakov, 1972). The fifth transgression - known as Valkatlen in Chukotka, Pelukian in Alaska, Attarman in Kamchatka - left a terrace with a height varying from 6 - 10 m to 20 - 25 m. An associated Arcto - Boreal molluscan assemblage contains also Boreal species, such as Mytilus edulis, Astarte borealis, Neptunea vinosa, Buccinum baeri (Ivanov, 1983; Zamoruev and Petrov, 1984). U-series ages of the Pelukian beds center around 100 ka (Hopkins, 1973). A continental equivalent of this thermochron, represented by the Konnergino or Elveneiveem beds, yields a 14C dating in excess of 60 ka (Biske, 1982, pp. 9 - 12), along with a number of finite, apparently, underestimates. Ca/Mg ratios show water temperatures varying from 3.8 to 13°C and 3 t o 7"C, respectively, for the Valkatlen beds of Chukotka and for the Pelukian beds (Svitoch et al., 1980), i.e. a maximal temperature value for the Pleistocene. The above data are consistent with the results of pollen analysis of the Konnergino terrestrial beds. In Konnergino time the Mayn River basin, 65"N,

161

presently occupied by a forest -tundra, was covered by taiga; this fact is also confirmed by fauna including the squirrel Tamiasciurus, bank vole Clethrionomys, timber beetle Hylobius albosparsus, carpenter ant Camponotus herculeanus, and the like (Svitoch et al., 1980). The above arguments unambiquously point that the Valkatlen - Konnergino thermochron is correlative to isotopic substage 5e. Some data suggest that another terrace-like feature - the Amguema, 5 to 20 m high, correlative, according to Petrov (1966), to the Karginsky terrace - occurs on the elevated shorelines of Chukotka. However, 14C datings yielded an Early Holocene age, falling in the range of 11 - 6.7 ka, for the Amguema terrace (Svitoch et al., 1980). Analogously, for the Bootlegger Cove clay in Cook Inlet, previously associated with the infra-Wisconsin ( = Woronzoff) transgression (Karlstrom, 1961, 1964, 1968), a l4C age range of 11 to 4 ka was reported later (Hopkins, 1973).6 The age of marine sediments of a 20 m high terrace at Point Barrow, northern Alaska, with an Arctic fauna and 14C datings in the range of 24 - 40 ka is still uncertain. Hopkins believes that the terrace occurs within a neotectonic high and represents a local (but not eustatic) feature. Hence, the Pleistocene climatostratigraphic succession on the Bering Sea coast appears t o be almost identical to that in the mountains of the north-east USSR. It is more difficult to reveal effects of climatic changes in the sections of the Yana - Kolyma plain, which represented a peculiar Arctic periglacial zone in Pleistocene time. The recognition of interglacials and glacials was initially considered unreasonable for the Yana - Kolyma and Yakutia plains, based on the assumption that the growth of frost an ice wedges is a continuous process. In fact, drastic diurnal temperature variations (no less than 19", e.g. - 24" to - 43") continue to result in the generation of frost cracks. In spring the cracks are filled with water which immediately turns to ice wedges. In virtue of the fact that the change in volume of water is an order of magnitude greater than that of dry ground upon freezing and thawing, the wedges once started would continue to develop into frost cracks in winter. Previously there may have developed in this way reformed ice wedges, vertically stratified, 6 - 7 m deep or, if accompanied by rapid sedimentation and ice generation, 30 - 40 m deep, with the width of ice veins measuring 8 - 10 m. The mean depth of elementary ice veinlets, 1 cm wide, reaches 5 m in such wedges. In 1952 V.N. Dostovalov showed that the age of reformed ice wedges can be determined from the number of elementary ice veinlets (Shumilov, 1982). The growth of separate ice wedges over the Yana-Kolyma plain was found to continue for no less than 12 ka and, hence, in duration of formation they can be well correlative with major stages of continental glaciation. Investigations carried out on the Yana - Kolyma plain in 1960 - 70s (Lavrushin, 1963; Popov, 1967; Cryogeological processes. . ., 1982) showed a well-defined stratification developed in systems of ice wedge polygons and, in the broader sense, "ice complexes", including ice lenses and reformed ice wedges, whose volume ac-

In the author's opinion, radiocarbon datings obtained on shells cannot be accepted as valid if they were not confirmed by estimates obtained from some other independent measurement.

162

counts for 60-90% of the rock volume. Wherever they occur, the systems of ice wedge polygons and the “ice complexes” are interbedded with less icy rocks represented by sediments of lacustrine genesis, which contain freshwater molluscs, peat bogs and buried soils. It is evident that phases of ice accumulation (“subsurface glaciation”) periodically alternated with thermokarst phases (i ‘deglaciation”), both types of phases being similar in intensity and areal distribution. However, correlation of subsurface glaciation and deglaciation phases with glacials and interglacials of the Atlantic sector is still debatable due to series of rejuvenated 14C datings. The genesis of sediments making up ice or “ e d ~ m a ”complexes ~ is a subject of controversy as well. Popov (1967) developed a hypothesis of their alluvial-lake origin, which is still popular. An alternative concept of the periglacial-loessal origin of ice complexes is under way (Tomirdiaro et al., 1983, 1985). Similar opinions were reported by American workers. The author (Zubakov, 1965) believes than an ice complex is a component of polygenetic congelifluction-sors assemblage, in which alluvial lake-pool and eolian facies replace each other along the strike and across the section. The scheme of Pleistocene stratigraphic subdivision for the Yana - Kolyma plain (Table 5.4) is a synthesis based on the data presented by Ivanov and Yashin (1939), Lavrushin (1963), Kaplina and co-workers (1978, 1982, 1983), Kaplina (19811, Tomirdiaro and co-workers (1983, 1985), Virina and co-workers (1984), as well as on materials of the Interdepartmental Meeting held at Magadan in 1982 (Biske, 1982). The scheme is climato-biostratigraphic being based, on the one hand, on an evolutionary succession of mammalian assemblages, established by Sher, Agadjanyan, and Zazhigin, and, on the other hand, on ice horizons interbedded with sediments of thermokarst lakes, spore-pollen zones, and faunal assemblages of insects (Kisilev, 1981). Exposures along the rivers: Krestovka, tributary to the Kolyma, 68”N; Bolshaya Chukochiya, 156”E, 70”N; Indigirka, 69” - 71”N; Alazeya; Khroma; and Adycha, tributary to the Yana, were used as type sections. According to new data reported by Virina and co-workers (1984), the Oler formation embraces a time interval from about l .2 - l .3 Ma to 0.4 Ma BP. The formation is subdivided into four biozones (Table 5.4) with three lower biozones being correlative with the Matuyama Orthomagnethem. Hence, zone 111 containing a forestdwelling bank vole and having the thickest peat horizon is considered as equivalent to the Anvilian-Early Enmakay thermochron (isotopic stages 19 to 21). The Oler Formation contains 4 to 9 stages of ice wedge polygons. This roughly corresponds to the number of kryomer isotopic stages in the interval of 1.2- 1.3 to 0.4 Ma. Evolutionary position and age estimates imply that the Oler fauna is equivalent to the Irvingtonian fauna of North America and to the Taman and Tiraspol faunas of the Ukraine. The Oler formation is overlain by a thick sequence of sediments which are not yet subdivided on climatostratigraphic ground; the sediments contain fauna of the Akansk - Utka type with Mammuthus primigenius pavlovae and Dicrostonyx

’

According to Biske, “edoma” is a mispronounced word “andoma”, i.e. hill, knob, mound in the local tongue of pomory (Kaplina et al., 1983).

I63

simplicior. These beds are also known as the Maastakh, Allaikha, or Khroma Formation. In sections along the Chukochiya river, near Lake Yakutskoe (71.0°N), the base of the formation is composed o f lagoonal-marine deposits which were drilled by a borehole; they were studied by Arkhangelov and co-workers (Biske, 1982, Vol. 2, p. 16). The deposits can be coeval to the Yanrakinot transgression and flood-plain peat bogs of the Ulakhan-Sular formation in the Yana River basin; the formation overlies river bed sands with Mamniuthus trogontherii and Equus cf. mosbuchensis. The Upper Pleistocene of the Yana-Kolyma plain is exposed in sections at: Molotkovsky Kamen along the Maly Anyui river (Kaplina et al., 1981); Stepnoi Yar along the Indigirka river (Lavrushin, 1963; Tomirdiaro et al., 1983); Duvanny Yar along the Kolyma river (Kaplina et al., 1978); and on the right bank of the middle Khroma river (Kaplina et al., 1983). Kaplina and Gitterman established three thermomers and three kryomers in the sections (Biske, 1982, Vol. 2, p. 14- 15). The Achchagyi thermomer ( = lower Shanga?, Krest-Yuryakh?, Kyl-Bastakh?) is represented by lacustrine deposits with: abundant freshwater molluscs Pisidium, Valvata, Anisus, Radix, and others; remains of wood Betula sec. AIbae, B . see. Fruticosue, Larix gmelini; water plants Pomogeton perfoliatus L., P. natans L., Menyanthes trijoliata L. and the like; and remains of such forest plants as Rubus idaeus L., Rosa cf. acucularis Lindl., Fragaria sp. These remains and pollen spectra of a larch- birch taiga, which grew at 71”N, as well as insect fauna, including ground beetle Carabus meander Fisch, leaf beetle Phosphaga atrata L. and others (Kaplina et al., 1983), all suggest an essential warming and northerly shift of geographical zones by 3 - 4 ” . An intra-Wurm (“Karginsky”) age of the horizon under study is suggested by a series of finite I4C dating falling in the range of 32 - 46 ka. However, such data as (i) effects of climate even warmer than the Holocene optimum; (ii) mass occurrences in the Achchagyi formation of Dicrosfonyx simplicior remains, which is a guide fossil for the Middle Pleistocene (Kaplina et al., 1983), and Marnmuthus primigenius pavlovae (Tomirdiaro et al., 1985); (iii) excessive, along with finite, datings above 52 ka and even 60 ka (Biske, 1982, pp. 10 and 36) obtained for the horizon, as well as extreme ages, above 46 ka for the overlying ice complex (Biske, 1982, Vol. 1, pp. 10 and 36; Vol. 2, p. 15) all testify against the intra-Wurm age. The above data ailow correlation of the thermochron under consideration with isotopic substage 5 e . The Oiyagos kryomer is represented by an ice complex with thick syngenetic ice wedges formed at extremely low winter temperatures, which dropped, according to Kaplina and co-workers (1981, 1983) down to - 70°, and possibly, even to - 100°C. Excessive 14C ages above 46 k a for the Woronzof edoma (Kaplina, 1981) and above 46.36 ka for the Oiyagos edoma (Tomirdiaro et al., 1985) suggest a correlation of the Oiyagos kryomer with isotopic stage 4. At that time the whole Laptev Sea shelf was drained, as suggested by a great number of ice wedge polygons on the Laptev Sea floor. The intra-edoma, so-called “Molotkovsky horizon” (Tomirdiaro, 1985), which consists, according to Kaplina and Lozhkin (Biske, 1982; Vol. 2, pp. 35-37), of two thermomers, is synchronous to isotopic stage 3 . The climatic conditions of the lower - Khomus-Yuryakh - thermomer resemble our own, while those of the up-

164

per - Kurenakh-Salan - thermomer were rather severe. These warmings were separated by the Kirgilyakh kryomer, a short, but well-defined cooling, synchronous to the Zhigansk (?) glacial stage in the Verkhoyansky Range area (Zubakov, 1974). Frozen corpses of different animals, named after the places of occurrence, such as mammoths (Berezovka with I4C datings at 44 s 3.5 ka; Shandrin, at 41.74 +- 1.29 ka; Berelyekh “Dima”, around 39.57 f 0.87 ka) and horses (San, around 38.59 k 0.112 ka), and others are almost solely associated with the Khomus-Yuryakh beds. This may be explained from the fact, that, although during this interstadial warming the “mammoth kryosteppe” was affected by thermokarst processes and bog formation, it still dominated the landscape. Steppes coexisted with bogs, and short-term warmings used to give way to short-term coolings. Therefore, natural death of animals under sharply differentiated natural conditions took place often, and conditions were particularly favourable for burial. Excessive 14C ages (above 36 and 45 ka) on plant remains from cores collected from Sannikov Strait at sub-bottom depths of 28 m and 32 m suggest that during the Molotkovsky thermochron the New Siberian Islands formed a part of the land and, hence, the “Karginsky” transgression could not take place there. The Mus-Khaya kryomer is characterized by (i) dominant micro-streaky semilaminated kryostructure; (ii) high gravity of ice lenses; (iii) narrow syngenetic ice wedge polygons; (iv) abundant azonal buried soils; and (v) abundant grass rootlets buried in-situ, in vertical position (Tomirdiaro and Chernenky, 1985). Pollen spectra and ecological analysis of the Mus-Khaya mammalian fauna dominated by horse, saigak, rhinoceros, bison and mammoth point, according to Garutt and Vereshchagin (Geochronofogy . . ., 1984), to the steppe character of the MusKhaya biome. At that time the relative role of eolian processes in the formation of ice complexes seems to have been maximal. The conclusion of Tomirdiaro and coworkers (1983) concerning a loess-Arctic genesis of the edoma during the MusKhaya kryomer is also confirmed by data of Kolpakov on the wide distribution of deflation sand deserts with sandblasted pebbles in the Lena River basin (Shumilov, 1982). Thus, the aridization of climate of north-eastern Asia and Alaska reached its peak in Mus-Khaya - Neptown time.

5.5. North America The subdivision of the glaciated Pleistocene of North America is based on lithostratigraphic data, i.e. on tracing till horizons and weathering crusts - “gumbotils” . Palynological analysis was poorly applied, and, hence, of the three interglacials, namely the Afton, Yarmouth, and Sangamon, only the latter has provided a reliable spore-pollen characteristic. Zoostratigraphic studies were more extensive. Much attention was given to the description of mammalians, insects and reptilians in North America (Taylor and Hibbard, in The Quaternary of the United States, 1965; Reppenning and Fejfar, 1976; Schultz, 1977; among others). But, because the occurrences of the fauna are mainly related to lithostratigraphic units of the unglaciated area, the Pleistocene biostratigraphic schemes in North America provide much less climatostratigraphic evidence than those of Europe.

165

At the VII INQUA Congress (USA, 1965) American workers presented some detailed chronostratigraphic schemes for the Pleistocene of the Central United States (Reed and co-workers; Frye and Leonard; among others, in The Quaternary of the United States, 1965) where the Nebraskan glaciation was subdivided into two stages, and the Kansan and Illinoian glaciations were subdivided into three stages, i.e. up t o 16 units were recognized in the pre-Wisconsin interval. During the ensuing years the traditional North American glacial scale was “extended” from 1.O Ma to 2.0-2.5 Ma. For example, Richmond (1970) proposed to correlate the Wisconsin with the Wiirm and Riss of the Alps; the Illinoian with the Mindel; the Kansan with the Gunz; and the Nebraskan with the Danube. The work of Easterbrook and Boellstorff (1981) from the Nebraskan Geological Survey gave unexpected impetus to the reassessment of the traditional Central North American Pleistocene sequence. Fission-track dating of five or six regional ash horizons and paleomagnetic studies of stratotypes designated as traditional glacial and interglacial “stages” showed that in different states of the USA tills with ages between 2.8 and 0.7 Ma were assigned to the Nebraskan; soils between 2.0 and 0.6 Ma to the Aftonian; tills between 1.7 and 0.8 Ma to the Kansan; and soils between 1.3 and 0.5 Ma to the Yarmouth. This suggests the necessity of revising the North American Pleistocene scheme. A complete revision of a section in the type area of the contiguous states of Nebraska, Iowa, and South Dakota allowed Easterbrook and Boellstorff (1981) to compile the following chronostratigraphic scale for glacial deposits of the central western area (Table 5.5). A great deal of new data has been recently collected by American and Canadian workers participating in “The Quaternary Glaciations in the Northern Hemisphere Project of the 24th Session of IGC (Richmond, 1983) and in The Quaternary stratigraphy of Canada (Fulton, 1984). Table 5.6 presents new data in two composite columns for ice sheets of the

Table 5 . 5 . New chronostratigraphic scheme of glacial deposits for the states of Nebraska, Iowa and South Dakota Subdivisions after Easterbrook and Boellstorff (1981) _____ ~ Till A1 - N zone Volcanic ash, 600 ka old Volcanic ash, 710 ka old Till A2 - N zone Till A3 - N zone with r event Till A4 - R zone Till B - R zone Volcanic ash, 1.2 Ma old Volcanic ash, 2.2 Ma old Till C1 Till C2

~

~

Local lithostratigraphic units - traditional nomenclature (The Quaternary 01the United States, 1965) _ _ _ “Kansan till” Pearlette volcanic ash “0” Hartford volcanic ash Cedar Bluffs till Nickerson, Santee and Hartington tills “Nebraskan till”, type section Unnamed till Coleridge (David City) volcanic ash Unnamed volcanic ash Upper Elk Creek till Lower Elk Creek till

I66

Laurentide and Cordilleran regions. In compilation of the first column the author used the data of Easterbrook and Boellstorff (1981), Karrow (1984), and Fenton (1984); the study of the territory north of 55"N is based on the results reported by Feyling Hanssen (1976), Fillon and co-workers (1981), Andrews and co-workers (1984), and Vincent (1984); the information on Greenland was obtained from the works by Hjort (1981), and Feyling Hanssen and co-workers (1982). The subdivision

Table 5.6. Tentative climatostratigraphic scheme of North American Pleistocene -~ I -Rock) Mountains and OCT Laurentide glacial area Cordilleran ~

~~

North Ailaniic maiine terrace,

I --1

Porl Huron KM

1

2 l

:I

Pinedale

-

J

-_

Fraser

~~

( = kluane Russel)

KM Lake Erie. 14- 1 5

Regression

I~~__~____ __ -

Nisrouri. 16 - 22

~-

-mM 2

Machnabb TM

Plum Poinr. 23

~

35

Middle

Olympia T M Spelcothem. Th/U 28-65

1 7

Guildwood KM

Bull Lake glaciation

4

Sangamon - Osler - Mlwnalbl T M T h / U 120

2

,4

Wando terrace 5-8m Sankaly Head Cliff. Th/U 133 7

_

2 E

".' i

1 s1

i2

~~

g zx ;.?

! &

TM?

E

_

_

Regression

2

.

socastee terrace 11-I2m T h / U 187-240

Speleoihem Th/U275-320 I

Yarmouth - Redcliffe T M

1

7

Canepath II? Talbot T h N 300-580 D/L 315-580

0.

14/16

Cedar Ridge ( = Oiting?)

16-lR,

Washakie Point

___

-~

Kegrcszron

U accamuu , 23 - 27 m Zone ".4"

H e i l l IWO- 1400

\

_

~

I67

of Pleistocene deposits of the Cordilleran region is based on the data of Richmond (1970, 1983), Harmon and co-workers (1977), Smith and co-workers (1983), Fulton (1984), and Smith (1984). In addition, the author referred to the results presented in the monograph The Quaternary of the United States by a group of authors and edited by Wright and Frey (1965), and in Abstracts of the Moscow Session of INQUA (1982). A succession of tills and buried soils in Nebraska and lowa states, given in Table 5.6 with due regard for revision of Easterbrook and Boellstorff (1981) helps to infer the stratigraphy of the Lower and Middle Pleistocene of North America. A similar succession was established by Stalker, Wesgate, Fenton, and others in the Wellsch Valley, Medicine Hap, Wascana Creek and other sections in the Canadian prairie (Fenton, 1984). In particular, the Twin Cliffs section in the vicinity of Medicine Hat in southern Alberta Province contains, judging from the earlier data of Fcnton (1984), seven climatostratigraphic units (members VI - XIV after Stalker), corresponding t o the Kansan, Yarmouthian, and tripartite Illinoian of Frye and Leonard (1965). The validity of such a minute subdivision of the lllinoian stage is supported by U-series dates on speleothems from caves situated in the Rocky Mountains and Mackenzie Mountains (Harmon et al., 1977; Gascoyne et al., 1981). Sections in the Great Lakes region, described in some detail by Dreimanis and Goldthwait (1973), Dreimanis and Raukas (1975), Dreimanis (1977), Karrow and Warner (1984) and others, have been designated as the type sections for climatostratigraphic subdivision of the Upper Pleistocene of North America (Fig. 5.9). The above authors recognized up to 16 climatostratigraphic units (six of them occuring in the Late glaciation interval of 13 to 10 ka) in the Upper Pleistocene of the Great Lakes region. Local schemes for the subdivision of Late Pleistocene on Baffin Island, in the Hudson Bay area and on Banks Island have recently been developed (Andrews et al., 1984), which can be compared in thoroughness to those for the Great Lakes region. However, 14C, Th/U and amino acid ages of newly recognized units in northern Canada have become substantially older in recent years. It is of interest that the last glaciation, as accepted in the Great Lakes Basin (Karrow, in The Quaternary Stratigraphy of Canada, Fulton, ed., 1984) and in northern Canada (Andrews and Miller, and Vincent, ibid.) does not coincide in extent; they are adequate, respectively, to isotopic stages 2 and 4, and stage 2 to substage 5d. A substantial uncertainty with respect to ages of numerous local lithostratigraphic units of the glacial and especially marine (Baffin Island and the Atlantic coast of the United States) Pleistocene still does not permit their reliable relation to the oxygen isotope scale. The correlation adopted by the author in Table 5.6 is thus tentative. The table shows that the traditional lower Pleistocene boundary - at the base of the Nebraskan ( = Shervin) till - is drawn in North America at the same level as in Europe and Siberia, i.e. at 1.O - 1.2 Ma interval, synchronous to isotopic stages 26 to 30. The numbers of glaciations established in North America and in Eurasia are equal. During all of the five glaciations the extent of ice sheets appears to have been correlative. As in Europe, the maximal advance of ice sheets fell on glaciations of different ages in different geographical regions. The difference in the time of

168

culmination of the Wisconsin ice sheet is particularly instructive. In northern Canada and in Greenland the maximal advance of ice occurred, according to Hjort (1981), Andrews and Miller (1984), and Vincent (1984), during isotopic substage 5d, 115- 110 ka BP, whilst in the Great Lakes region the maximal - Late Wisconsin - glaciation took place 22- 18 ka BP, during isotopic stage 2 (Dreimanis and

E A R L Y 00

SCIOTO S U B L O ~ ES U B L O B E

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LACUSTR.0EE (OIl0,N.Y.

NOTE: PROBABLE

LOBES

INDICATED

BY

THEIR

FIRST

LETTER

Fig. 5.9. Proposed correlation of litho-climato- and biostratigraphic units of Early and Middle Wisconsin age in the Great Lakes region (from Dreimanis and Goldthwait, 1973, fig. 4).

169

Raukas, 1975; The Quaternary Stratigraphy of Canada, Fulton, ed., 1984). The fact of repeated (up to 8 - 10 phases) formations of ice sheets which spread southwards t o 40”N itself suggests an unusually great amplitude of climatic changes in North America during the last 1 Ma. However, the paucity of palynological information and inadequate development of regional and inter-regional stratigraphic correlation make quantitative paleoclimatic reconstructions difficult. The most reliable reconstructions of landscapes and climate of the last ice age were presented by Washburn (1979 - 1980) and by a group of authors in the monograph Late Quaternary environment of the United States (Wright and Porter, eds., 1982). The above authors believe that in Late Wisconsinian time the tundra zone shifted 1500 - 2000 km south. At 35 - 40”N the then mean annual temperature is estimated to have been lowered by 8 - 10°C. A similar situation is suggested for the earlier glaciations. The reconstruction of climate occuring in warm intervals poses a more severe problem. Thus, the study of freshwater molluscs led Taylor (1965) to the conclusion that “All of the pre-Wisconsin climates were more maritime than the modern climate at the sites where the fossils were collected. None of the pre-Wisconsin faunas known could live in the present combination of hot summers and bitterly cold spells during the winter . . . Pre-Wisconsin climates, both glacial and interglacial . . . differed among themselves mainly in amount and effectiveness of mean annual precipitation, and summer temperatures” (The Quaternary of the United States, 1965, p. 603). Hibbard and co-workers who analyzed the composition of a mammalian assemblage came to the same conclusion. They write: “Present-day climates, with their seasonal extremes of temperature and aridity, are . . . atypical , . . of the Pleistocene” (ibid., p. 515). Thus, the Cudahy fauna of “Late Kansan” age which lived in the region between 37 and 30”N, was made of musk-ox, collared lemming, tundra bog lemming, vole, shrew and other animals of the boreal zone, as well as of jumping mouse Zapus sandersi, horse, emperor elephant. The interglacial Borchers (Yarmouth? time) and Cragin Quarry (Sangamon) faunas include warmth-loving animals, such as turtles Geochelone dwelling in non-freezing reservoirs, cotton rat Sygmodon hilli, and the like. Frey ( 1 965) believes that “climatic conditions gradually deteriorated after Miocene time, resulting at the end of the Pliocene (Early Nebraskan? V.Z.) in conditions probably more adverse to plants and animals than at any other time in the Cenozoic . . . At the beginning of the Pleistocene there was a marked climatic improvement, culminating in Kansan time” (?V.Z.)* (ibid., 1965, p. 621). Unlike Taylor, Frey thinks that “during the glacial ages the climate was much more severe than at present” (ibid). The above quotations show that the objective paleoclimatic reconstructions for the territory of the United States have not been precisely substantiated by stratigraphic and paleobioiogical evidence. In this connection of particular interest are the information of Harmon and coworkers (1977) on Pleistocene chronology and paleo-temperatures obtained through analysis of the oxygen-isotope composition of calcareous sinters of speleothems and their dating from 230U/234U ratios. The above scientists studied 36 speieothems

’ It is not clear from the quotation what “Kansan

time” implies.

170

from caves located in the Rocky Mountains and Mackenzie Mountains in northern Canada, 65"N; they established age groups for four time intervals, namely 90- 150 ka, 185 -234 ka, 275-320 ka, and over 350 ka. Because stalagmites cannot build up at a temperature below O", the periods of their growth may be considered as obvious thermochrons. Data on speleothems are in good agreement with the oxygenisotope scale. During the years which followed, an additional 30 datings on speleothems were obtained from caves in the Central Western United States, Virginia and the Bermuda with concomitant paleotemperatures (Harmon and Schwarz, 1981). Table 5.7 shows that in the zone affected by the Gulf Stream the mean annual air temperature differed from that at present by 8 - 13°C and 6 - 4°C in glacial and interglacial times, respectively. The recognition of two thermochrons within Illinoian time is also supported by new data on marine terraces off the Pacific Coast where Wahrhaftig and Birman (1 965) suggest the presence of a series of 13 - 20 marine terraces up to 400 - 450 m

Table 5.7. Climatochronologic series on speleothems from caves of North America. Ages from U-series measurements, in ka; A T("C) from 'H/"O measurements (Harmon and Schwarcz, 1981) _______ OCT Bermuda Western Virginia Kentucky ~____________ Age (ka) AT("C) Age (ka) AT(T) Age (ka) AT('C) __ ________ 2 25 -1 29 - 13 ~ - _ _ - - _ _ _ __--_ _______ 4 60 -5 _____~___ -_____ ~5a-c 97 +I 99 -7 101 -3 I04 -8 104 +8 5d

109 112 115 118

-5 -7 -4 - 10

5c

~

I72 176 185

____

+8 +6 -2

-2

112

-1

122

0

127 159 169

-3 0

156

-8 -3

172 174

+6 +4

171

+4

195 217

+2 +2

6

7

106

____ -___

_____.

______

171

high. The ages of marine molluscan shells from three lower terraces in San Diego, determined from amino acid content, are about 120, 200, and 300 ka (Karrow and Bada, 1980), i.e. they coincide with isotopic stages 5, 7, and 9. Higher terraces are believed to correspond to isotopic stage 11 and further. A series of 20 marine terraces is developed on Saint Clemente Island off California. U-series datings on corals A//oporacalifornica and shells of molluscs Epilucina and Tegula have fallen in the age ranges of 80- 105 and 120- 127 ka for terraces I and 11, respectively, while amino acid (D/L) ages of terrace V vary from 415 to 575 ka (Mush and Szabo, 1982). This suggests that the terrace succession there adequately reflects all the eustatic changes in sea level throughout the Pleistocene. A more complex and ambiguous pattern of glacio-eustatic sea level variations was established on the eastern coast of the United States. There terraces are poorly expressed in topography and their correlation is obscured by neotectonic movements. Microfossil analysis permits recognition of three zones on the basis of ostracods in the Pleistocene of North and South Dakotas, namely zone A, corresponding to reversed polarity formations: the Wicomico ( = Penholloway) and the Waccamaw ( = Croaten, James City), which make up a level 26 - 30 m high; zone B, corresponding to normal polarity sediments of terrace level 9 - 18 m high (Talbot, Canepatch); and zone C , corresponding to a level of terraces of Princess Anne, Socastee 1 to 10 m high (Cronin, 1980). U-series (McCartan et al., 1982) an amino acid (Wenmiller and Belknap, 1982) datings on corals and shells revealed a more complex relationship between marine sequences from place to place. However, Useries and D/L datings are not always in harmony. Since the U-series technique is more advanced, four or five transgressions can be established from the results obtained (McCartan et al., 1982). D/L and He/H datings show that the oldest - Waccamaw - transgression yielding such index fossils as Mercenaria permagna (Conrad) and Conus wuccamawensis B . Sm., which left a terrace 26-30 m high, took place about 800 ka and 1.O - 1.4 Ma BP, respectively. The transgression probably corresponds to the pre-Nebraskan thermochron with fission track ages centered around 1.27 Ma (Easterbrook and Boellstorff, 1981). The second transgression with U-series and D/L ages, respectively, around 760 and 650- 780 ka are ascertained tentatively and correlated in Table 5.6 with an interglacial which separated the Nebraskan and Kansan Glaciations. U-series ages of the third (Canepatch - Talbot) transgression containing M . mercenaria and Argopecren solariodes (Heil.) are in the range of 300- 580 ka with an average of 460 f 90 ka; the transgression is presumably correlative with isotopic stages 15 to 11. U-series ages of the fourth marine terrace, 9 - 12 m high, containing M . campechiensis Gmel. and Dosinia elegans Conrad range from 187 to 240 ka; the terrace probably corresponds to the infra-lllinoian Interglacial which has not yet been established in North America; the interglacial has been also recorded from speieothems. And, at last, the youngest transgression (Wando), dated within the Sankaty Head section, Massachusetts, by U-series on corals at 133 k 70 ka and by D/L at 120- 140 ka, by far corresponds to the Sangamon Interglacial. Similar age estimates for the t w o last transgressions were also obtained on corals from the Bermudas (Harmon et al., 1983).

172

The above data suggest that a succession of Pleistocene climatic events in North America appears t o have been more complex than accepted earlier. During the last 1 Ma at least five, instead of four, previously adopted glaciations occurred there. Of particular significance for paleoclimatic reconstructions is the information on pluvial events in desert intermountain depressions of the Great Basin and of the south-west of the United States. At present this area is occupied by small salt lakes and dry beds (playa). In the past there were about 120 freshwater lakes (Bonneville, Russell - Mono, Searles, Manly in Death Valley, and others) with water level exceeding the present one by 50- 100 m or even 200- 300 m. The largest of the Lakes, Bonneville, about 20,000 sq.km, is believed to have overflowed into the Pacific Ocean via the Snake River. Studies of the Late Pleistocene history of the lakes were performed by Flint (1957, 1971), Frey (1965), Morrison (1968). Interesting new data covering the whole of the Pleistocene and the Late Pliocene were reported by Eardley and co-workers (1973) and Smith (1983, 1984). Drilling in Lake Bonneville (the remainder of which is the Great Salt Lake), Utah, and Lake Searles in south California, 36"N, 117"W, showed that the sections drilled provide records of all climatic changes occuring in the arid zone of North America. Thus, the Lake Bonneville section gave evidence of 28 climatic cycles for the last 800 ka. Pluvial periods are represented by lake mud, while interpluvial phases are composed of evaporites (mainly carbonates and sodium chlorides) and buried salt soils. Estimates reported by different authors indicate that during the pluvial periods the mean annual and July temperatures were, respectively, 2.5 - 5°C and 4 - 5°C lower than at present. Precipitation, which at present varies from 100 mm in the south of the region to 300 mm in the north, then exceeded the present values by 70- 100% and reached 400- 500 mm (Morrison, 1968). Taking into account that most pluvial lakes in the Great Basin did not have glaciers within their drainage areas, the pluvials could not be attributed to melting of the mountain glaciers. American geologists believe that the high level in pluvial lakes was due t o a decrease in evaporation caused by a decrease in temperature, increase in cloudiness and precipitation, i.e. shift of the zone of cyclonic activity south from its present position approximately 15" by in latitude (Flint, 1957, 1971; Morrison, 1968). Estimates of Smith and co-workers (1983) indicate that a Wisconsin decrease in annual temperature by 10°C ensures an eight fold increase of run-off in the Lake Searles basin.

5.6. The Arctic and sub-hrctic

Reconstructions of the past Arctic climates are much more complicated than of any other region of the Earth. In fact, as was stated by Saks (1963, p. 90): " . . . fauna of the present-day North Siberian Seas is the most kryospheric of all the known faunas, it seems impossible to imagine any other fauna which would require lower temperatures for normal development, since seawater freezes at minus 1.8"C. That is why the fauna of these seas would be indicative of only those periods when temperatures were higher than the present." Glacial marine deposits are also forming now, which makes it very difficult to find lithological differences between them and the till. Saks (publications from 1945 to 1963) laid the foundation for a

I73

climatostratigraphic division of the Arctic Pleistocene. He proposed a scheme in 1948 based on the sections in the lower reaches of the Yenisey; later his scheme was applied to the whole of the Arctic. It includes four glaciations: ancient (?), maximum, Zyryanka and Sartan, as well as three interglacial transgressions: Northern, Boreal and Karginski. The scheme served as a certain stratigraphic Esperanto, being the Arctic equivalent of the Alpine scheme. However, its disagreement with local features even at the Yenisey section were already established by 1957. Three main tendencies can be revealed in the later works: (1) more complicated scheme with a revision of unit genesis; (2) the lower units of the scheme were recognized as older (the Pliocene); while (3) the upper units of the scheme were considered younger on the basis of radiocarbon datings. The complications started from the changes in the genesis of the middle part of the succession of Boreal transgression - i.e. Sanchugovka beds. (3.1. Lazukov (1970- 1972), Arkhipov and Yu. A. Lavrushin (in 1957- 1970) recognized Sanchugovka beds as being of glacial-marine origin, while Zubakov (1956 - 1968) referred them to shelf-glacial assemblage. Appropriately the Boreal transgression was divided into two parts by the TazIYenisey glaciation (see Zubakov, 1972, Lazukov, 1972). A number of investigators (Kaplyanskaya and Tarnogradsky, 1975, 1984), Astakhov (in Arkhipov et al., 1980) related Sanchugovka beds of Siberian rubbly loams to ground moraine believing that fossil fauna yielded by rubbly loams has been glacier-entrapped from the underlying surface, though nobody could claim the locality of those underlying marine beds. To overcome this ambiguity, a hypothesis was advanced: it assumes that the glacier moved towards the plains from the Kara Sea shelf rather than from the Siberian plateau and the Urals (as indicated by mineral and petrographic evidence). Thus the number of glacial (or glacial-marine) beds has increased t o six (Gudina, 1969; Zubakov, 1972, 1974) and interglacial beds yielding boreal fossil faunas are believed to number five. Another group of researchers, who are ardent adherents of “glaciological” hypothesis emphasizing a complicated structure of the Pleistocene succession (the presence of erratic masses, glacial tectonics and the like) and its poor knowledge, claim that “reliable fixation of interglacial beds in the Northern Pleistocene section will not become possible in the near future” (Astakov, 1984). The third group of researchers (Zagorskaya et al., 1965; Suzdalsky, 1974; Zarkhidze, 1981; Danilov et al., 1983) dispute the glaciation of northern Eurasian plains attributing the lower part of the 200 m marine section of the ancient buried valleys to the Late Miocene/Early Pliocene. The widespread use of radiocarbon dating has led to a new tendency of relating deposits, earlier believed to belong to the Sanchugovka, Kazantsevo and Zyryanka stages, to the Karginski and Sartan stages (Kind, 1974; Andreeva et al., 1981; Arkhipov et al., 1977; Svitoch et al., 1980). These three tendencies have their proponents among western scientists. This is illustrated by the change in the age of marine deposits of the Baffin Island in the type locality of the Clyde Foreland formation. Earlier all 12 units of this suite were I4C-

174

dated as belonging to the Upper Pleistocene (Feyling-Hanssen, 1976). On the basis of U-series and amino-acid analyses their age was determined as older (Andrews et al., 1983, 1984) and sometimes much older. Thus, for instance, foram-zone Cassidulina teretis initially attributed to the St Pierre Interstadial and correlated with Jameson Land beds in Ladin Elv formation (East Greenland) was recognized as the Lower Pleistocene on obtaining D/L values equal to one million years (Feyling-Hanssen, 1982). Considering these contradicting tendencies the climatostratigraphy of the Arctic Pleistocene is in an extremely debatable state now, some authors even seem to deny Table 5.8. Climato5tratigraphic units of the Artic and sub-Arctic Pleistocene OCr

1

North American A r c l ~and Sub~Arutlc and Spitrbergen

S o ~ i e iW e w r n and Central Arctic and Sub.Arctlc

I

1

Mountain glaciation: No r i l k a - Sopkri KM -

Ballin Land - Russell KM "A" lone (Clyde Foreland i 1

E A

E I

Iron Strand'! KM. TL 29 - SO? c n u u

~

c

lgarta

____

(=

'.Kaigin53, Th/U 8 0 - 114

[Yenisei ( = T u ? - "Zyranka I"? - Murukhta?) Kb

6

SalekhardRogovaya glacial and ice-shelve asiemblage

7

~

8 - 10

u 3

'

"Karantrevo - Pupkovo"- Timan - Rodionov T M Cyprinu islandicu - Unro hybridu T h / U 164- 2331 -

.m 3

3-

I r.

2

~

~

Merso - Makarikha sands and gravel Monimurhus cf. Lrogonlherri (KM?)

I2 -.

~

5

13

Clandulinu nipponico zone

14

IS

0

2

16- 18

3

-

Elphrdwn - Islondrellu - Alubimmotdes zone

5, 0

19-21

m

F

~

22 - 24 ~

25?

'%

Elph. suborrfrcum

D

-

__

~

?

E:

2 A

~

20?

2

Proloslph. orbrczilore Elph. ~ ~ c ' ~ ~ u rubzone rurn

2

Basal beds of canyons sands, gravel, clay with Cvrerrsso lacusrris - Cyclorello boicolensis

8

-

Cossrdulrno subucuro rubzone Torellkjegla TM. TL 413 t 62

g

~

-

Crrssrdulmo ierel!s subzone ("G") Th/U 2 150, D/L IWO?

Bakhla - Salemal'? - Nartset? KM

Padimei - Kochos Yakovleva formation (TM?) Cvrroduna anguslo Milir,line/lo Emndrs ~

II

Clyde Foreland 1. Driir II

Is1 Boreal transgression-

? '

Elph. excuvorum wbzone

-

Cuss. rereris

Clyde Foreland 1'. Drifr IV Bankstill

rl

5 g

Protoelph orbiculare Coss. rereits subzone, Morgan Blulis T M

5

Clyde Foreland 1. Drift V Duck Hawk Bluiit KM

1 $2

I

I75

its possibility. However, as Table 5.8 shows the problem is not so hopeless. The table is based on numerous recent publications, just to name the most important: Gudina (1969), Gudina et al. (1983, 1984), Zarkhidze (1981), Danilov et al. (1983), Arkhipov et al. (1977), Lazukov (1970- 1972), Levchuk (1982), Andreeva et al. (1981), Troitsky (1975, 1979), Arslanov et al. (1980, 1981a,b, 1983), Guslitser and Izaychiev (1983), Makeev et al. (1981), Avdalovich and Bidzhiyev (1984), Vasilchuk et al. (1984, 1985); Feyling-Hanssen (1976, 1982), Fillon et al. (1981), Hjort (1981), Vencent (1984), Hopkins (1965, 1979), Herman and Hopkins (1980), Clark (1982), The Quaternary stratigraphy of Canada . . . (Fulton, Ed. 1984). The table also includes new materials of Andersen et at. (1983), Mangerud et al. (l981), Miller et al. (1983), Lindner et al. (1983), Ruddiman and Mclntyre (1981), Boulton et al. (1985) and many others for Norway, Spitsbergen and the Norwegian Basin. Comparing the materials of the cited authors with his own observations on the section in the lower reaches of the Oh and Yenisey (Zubakov, 1972a,b; Geochronology of the USSR, Zubakov, Ed. 1974) the author divides the whole succession into 17 - 20 climatostratigraphic units, thermomers, the fossiliferous beds, being the most important ones. Organic remains indicate organisms that now inhabit more southern areas or are completely extinct. These are foraminifera, ostracod, marine mollusc assemblages requiring positive bottom water temperatures, remains of arboreal plants, as well as taiga and tundra - forest sporepotlen spectra. The kryomers are not so well distinguished, being represented by

Recent

-

3

4-5d

5e

7

8- 10 al . - I

13-15

16-18 19-a 22-24

Fig. 5.10. Sequence of continental glaciations and marine transgressions o n the North A5ia coati (compiles by the author from data of Gudina et al., 1979, 1983; Zarkhidze, 1981; Danilov er al., 1983; Zubakov, 1972, 1974). I - OCT, 11 - alluvial sedimentation, 111 - glacial area, I V marine area. ~

176

glacial and glacial - marine lithological complexes and by beds yielding tundra and tundra - steppe spore-pollen spectra. The author believes that the upper limit of occurrence of Elphidiella oregonense Cush. et Gir. is the boundary of the Arctic Pleistocene. According to Voorthuysen and Doppert its range zone coincides with the Icenian base in the Netherland section, i.e. being 2.4-2.2 Ma old (see Zagwijn, 1974). In the Bering Sea area it characterizes the Anvilian succession the top of which (by Hopkins et al., 1974) corresponds to the Jaramilio event. As shown by Gudina et al., (1984) Elphidiella quasioregonensis Cud. in the Ayon Island section (see 5.4) does not spread higher than the BrunhedMatuyama transition. The Icenian formation is topped by the Menapien kryomer, according to West, its age being 1.2 - 1.1 Ma, Zagwijn (1974), though, claims that it is the Eburonian kryomer that crowns it (1.4- 1.3 Ma). Its Icelandic equivalent is the Breidavik formation overlapped by reversed-polarity lavas, K/Ar dated as 1.12 0.9 Ma old. Summing up, we can determine on the basis of this not very definite evidence the age of the top of E.oregonense range zone in the Arctic as 1 .O - 0.9 Ma. This would also limit the age of the bases of Bolshaya Kheta and Kolva formations of buried valleys in Northern Siberia and the Pechora Basin as well as the range zone of Cibicides grossa on Baffin Island (where no Eoregonense was found). Thus the erosion phase caused by the development of buried valleys on the Northern Eurasian shelf would correspond to the “B” nonconformity, its age being also determined as 0.9 Ma (Zagwijn and Doppert, 1978). It thus follows that the oldest kryomer of the Kolva formation and Clyde-Foreland formation are likely to be equivalent to isotopic stage 22 - 24, that is 900 f 50 ka. (Fig. 5.11). Thus the sections of the Bolshaya Kheta and Kolva formations seem to encompass two large climatic-sedimentary cycles, corresponding isotopic stages 20(22?) - 19 and 18- 15. The Bolgokhtokh and Ust’solenaya transgressions are believed to commence in the late glacial environment as indicated by Iithological and faunistic evidence. This would lead to the inference that the bottom water temperatures were negative at the beginning of the transgressions (Danilov et al., 1983; Gudina and Khoreva, 1984). The second half of these transgressions developed in a more favourable climatic environment. This is indicated by ostracods with Elofsonella concinna (Jones), Eucytherideis punctullata (Br) and others signifying, as shown by O.M.Lev, a non-freezing sea with winter temperatures not lower than 5°C (Danilov et al., 1983, p. 105). Warm-water foraminifera comprise 65% in West Siberia and up to 80% in Pechora Basin during the Upper Kolva (the Ob and Ust’-Solenaya) cycle. Foraminifera assemblage includes also Myliolinella pyriformis (Schl.), Elphidiellaflorentinae Shypack ( = E. tumida Gud.), E. hannai Cushm. and others together with numerous Pacific Ocean migrants, such as Glandulina nipponica Asano ( = Tappanella arctica Gud.), Islandiella limbata (Cushm.). The Kolva - Bolshaya Kheta mollusc fauna also contains many Pacific Ocean species (Nucula tenuis Mont., Serripes groenlandicus (Brug.), Tridonta borealis (Shum.) and others. In most cases Atlantic species were absent except in ostracod fauna. Zarkhidze and Slobodin attribute this to a closer relationship of the Arctic Basin with the North Pacific than at present (Danilov et al., 1983). We can also suggest that the Barents Sea shelf during Kolva time (up to the 12th isotopic stage?) was

*

177 I 0 0

Num bcr of specimens i n l O O g sediment

0 0 0

-

0 0 m

-

0

m N N

Ln 0

I 0 N

I @ ,--+-- i 1 = . . i I.......

t t C o s s isduubl ioncau t o

__

1a I

Elphidiurn excavatum

1

I

I

I

I

1

I

I

I

# = -

I

I

Elphidiurn 0 1 biurnbi Iicaturn

111111....

I 11-11..

-

......*

Elphidiurn su b o r c t i c u r n I

I

Cibicides rotundatus

1

1

1

1

1

1

I

Rotalia columbi ensis

ZONES

I

zone

ISAl

zone

Fig. 5.1 1. Stratigraphy of the Clyde-Foreland Formation illustrated by the vertical distribution of 13 selected species of foraminifera. The average number of species per 200 g sample and average number of specimens per 100 g sediment are indicated to the right (after Feyiing-Hanssen, 1976, fig. 7 ) .

178

either a land-mass or a very shallow basin for the Gulf Stream; it would also indicate a higher Farae - Icelandic Sill. Bottom water temperatures, nevertheless, were higher than at present, ranging from - 1 to + 2”C, but the salinity was close to normal, that is 34- 34.5% along the whole length of the Siberian shelf, from Pay-Khoy to Chukotka (Gudina et al., 1984). The Kolva cycle terminated by a marine regression and formation of sandy-pebble beds (Messo Suite, according to Saks, 1948). Troitsky states (1979) that sandypebble beds contain remains tills. I f it is really so, then this glaciation would correspond to the Sarchikha kryochrone in Siberia and the Elster and Oka Ice Ages in Europe (Fig. 5.10). Important paleogeographic changes in the Arctic occurred in the Middle Pleistocene, manifested by a stronger relationship between the Arctic Basin and North Atlantic. It is testified by invasions of North Atlantic foraminiferas and molluscs (Elphidium atlanticurn, Melonis Zaandamae; Cyrtodaria angusta Nest. et West) reaching northward t o Severnaya Zemblya. The Padimey - Kochos transgression is most likely synchronous with the Holstinian. They both started in the late interglacial environment, which is well fixed in the sections near Hamburg (Grube, 1984) and also on the Yenisey where the lowest part of the “B” unit of the Sanguchovka formation (according to the classification by Troitsky, 1966) belongs to this time. The Sanchygovka basin is characterized by cold-water fauna species with Joldiella lenticula and Bathyarea glacialis. The upper part of the same unit yielded Sub-Arctic and boreal molluscs including Pecten islandicus (Mull), Buccinum undatum L. and others. Slobodin found here the richest foraminifera complex with Protoelphidium ustulatum and P. lenticulare (Zubakov, 1972a). This part of the transgression fixed on the Pechora River by a specific ostracod complex abundant in Cytheropteron and by the presence of molluscs Cyrtodaria angusta, seems to correspond to a thermochron. Biske and Devyatova (in Antropogene period in the Arctic . . ., 1965) pointed out that taiga with broad-leaved trees grew over the coasts of the Padymey Sea ( = Likhvin Sea). Thus, the data available indicate that the climate at the height of the Padymey - Sanchugovka transgression was warmer than it is at present. The Torellkjegla beds on Spitsbergen TL-dated as 413 k 62 ka (Lindner et al., 1983) and the middle part of the Cibicides grossa zone on Baffin Island are found to be equivalent t o the Padymey thermomer. The age of the overlying succession of boulder-loams of the Rogov formation of the Pechora River, of the Salekhard formation on the Ob, and of the Sanchugovka formation on the Yenisey is approximately determined by Mammuthus cf. trogontherii remains in the base of Rogovaya Beds on the Makarikha River found by Zarkhidze. As it was earlier established (Chapter 3) the last appearance of M . trogontherii occurred during OCT 12 (470-500 ka). Since it is difficult even on well-established sections in some cases to distinguish marine interglacial deposits, i.e. glacial-marine deposits and till developed out of marine deposits, the whole succession of boulder-loam is to be considered as a shelf-marine assemblage, this is an involved paragenesis of undivided genetic types and climatomers, corresponding to four isotopic stages, from the 13th to 8th (or, perhaps, from the 1lth to 6th). Undoubtedly, the uppermost part of this assemblage (unit “G”) in the lower reaches of the Yenisey is a continental moraine, as it was already pointed out by the author

179

(Zubakov, 1972, p. 27 and on other pages). Kaplayanskaya and Tarnogradsky (1975, 1984) confirmed it. The relief-forming sands with the broken fauna (Malyshevka, Us[’-Port, Nikitin, Vashchutka, Sabun Beds) overlap the moraine composing a fluvioglacial cover, formed at the time of the degradation of the Samarovo and Yenisey Glaciation. As shown by Saks (1984) and later updated by Troitsky (1966, 1975) Kazantsevo beds in the lower reaches of the Yenisey, yielding Arctic- Boreal molluscs with Cyprina islandica, Buccinum undatum and others are believed to be the traces of the first boreal transgression. In its type section they overlay Sanchugovka boulder loams and underlie Zyryanka boulder loams. While in other areas the deposits of the first Boreal transgression are found on some other stratigraphic situation, that is they intrude into the Sanchugovka formation and its equivalents (Salekhard formation on the Ob and Rogovaya formation on the Pechora river). These are the Pupkovo, Kharsoim and Vastyanski Kon’ beds. The Pupkovo beds are represented by marine clays with Boreal molluscs Cyprina islandica and Macoma baltica and a pollen diagram reflecting fir - spuce taiga underlying the peatbog with Lycopodium clavatum, Osmunda cinnamomea. The age of mollusc shells was dated (by U-series) as 233 t- 10 ka at R = 1.15 (Zubakov, 1972a, 1974) corresponding to the 7th isotopic stage. Rodionov beds, their continental equivalent, were described by the section near the Kipievo Settlement on the Pechora (66”N). The thermomer is represented there by a lacustrine - alluvial succession separating two tills. It is characterized by spore-pollen spectra of spruce forest with a small addition of broad-leaved species (Corylus, Quercus, Ulmus),the Carpoidea contain Ajuga reptans L. which at present grow some 400 km to the south. Rodionovsky beds yield shells of unionids (Unio tertius Bog’, U. cf. hybrida Bog.), at the top fossil rodents were found: Dicrostonyx ex. gr. simplicior Fejf., Lemrnus cf. sibiricus Kerr., Microtus sp. (Guslitser and Isaychiyev, 1983). In the marine sections of the Pechora basin Timan foraminifera complex and “the third mollusc complex” by Zarkhidze (Vastyanian) correspond to the Pupkovo - Rodionov thermomer as described by Baranovskaya. The “third mollusc complex” contains Boreal fauna in substantial quantities as found for the first time in the Pleistocene history; these are Spisuia eiliptica Br, Zirphaea crispara (L.), Modiolus modiolus L., Neptunea despecta L., Bussimum undatum and others (Danilov et al., 1983, p. 101). The Pupkovo - Vastyanian mollusc complex composition does not differ from that of the Kazantsevo complex in its type locality at the Lukovaya stream (the lower Yenisey), neither from that of the Kazantsevo complex on the Agapa river, from which shells of Cyprina islandica L. were U-series dated as 164 5 ka (Zubakov, 1972a, 1974). This allowed Troitsky (1979) to refer the Pupkovo and Vastyanski Kon’ beds to the Kazantsevo transgression and to correlate them to the Eemian. The U-series dates (233 - 164 ka) and Dicrostonyx sirnplicior fauna, however, do not agree with the fact that the described thermomer was correlated with the 5th isotopic stage, though the correlation of the Pupkovo beds with the Kazantsevo Beds seems to be correct. The Baffin Island sections (Fig. 5.1 1) display the Cassidulina teretis zone which is believed to be equivalent to the Pupkovo - Vastyansky Kon’ thermomer. It was initially correlated with the St Pierre Interstadial, while later it was referred to the

*

180

Earlier Pleistocene (Feyling-Hanssen, 1982) which is rather doubtful and even erroneous. In north-western Europe the beginning of the Boreal transgression is marked by Early Eemian and Early Ipswichian beds which Kukla (1977), Sutcliffe (1985) and Bowen (1978) correlated with the 7th (and the 9th) isotopic stage. Thus, older ages of the beginning of the Eemian - Boreal transgression (as well as the Thyrrenian and Karangatian) corresponding to the 7th isotopic stage should not be considered as a regional phenomenon or an accidental conclusion. The first Boreal transgression followed the maximum glaciation of the high latitudinal areas of the Northern Hemisphere (Samarovo - Saalian - Illinoian) which left the Wedel-Jarlsberg Land moraine on Spitsbergen, TL-dated as 313 k 47 to 229 f 34 ka (Lindner et al., 1983). It starts the third paleogeographic stage in the Arctic history, encompassing isotopic stages 7 - 6 - 5. The second Boreal transgression was distinguished by Saks (1948), who called it the Karginski stage. Troitsky (1966, 1975), though, first believed (later he changed his opinion) the sands of the Karginski section in its type locality on Karginsky Cape and the Kazantsevo section on its type locality at Lukovaya stream (both sections being in the same area, on different shores of the Yenisey) t o be synchronous. Troitsky (1966) referred the deposits on the 2nd level Yenisey terrace to the Karginsky stage as well, which were later 14C-dated to the period, ranging from 22 to 30 ka. This ambiguity in the nomenclature gives rise to numerous discussions and problems. A series of terraces 45 -60 and 80- 100 meters high on the Arctic islands (Siberian side) can be related to the second Boreal transgression. Similar terraces at the foothills of the Byrranga Mountains are found at 120- 200 meters. The height of the terrace at the Gydan and Yamal Peninsulas reaches 40-60 meters. It is overlapped by thick peat bogs, which were I4C dated as 32.7 - 40.7 ka (Avdalovich and Bidzhiyev, 1984). The Karginsky terrace on Novaya Sibir Island, not glaciated and subject to glacioisostatic movements, was 20- 35 meters high and yielded a mollusc complex with Astarte borealis var. placenta Morch, A . montaguif. typica (Dillw.) Yens, and others (Ivanov and Yashin, 1959). Foram complexes of the second Boreal transgression are called the Shchuchya complex in West Siberia and Ponoi complex on the Kola Peninsula (Gudina et al., 1983), with Cibicides in the Pechora Basin (Danilov et al., 1984,. p. 53) and on Baffin Island where it is the Zslandiella islandica zone, yielding also molluscs Chlamys islandica (Mull.), Astarte borealis (Chemn.) and others (Feyling-Hanssen, 1976). Serial 14C datings do not definitely indicate the age of the deposits of the second Boreal transgression. As shown by Andreyeva et al. (1981), out of 120 datings on Karginski beds in the northern West Siberian lowland (including even younger continental and liman sediments) 18 fell outside the range of 46 - 23 ka. A number of extreme and out-of-the-range datings were obtained on the so-called “Kharsoim beds”9 in the lower reaches of the Ob (Arkhipov et al., 1977). 14C datings of 28.2 -46.95 ka were determined for Cape Broughton beds of the Islandiella islandica zone (Feyling-Hanssen, 1976). The age of Valkatlen and Mga beds correlated well with the second Boreal transgression; it was I4C-dated as belonging to the The name is reserved (Zubakov, 1972a) for the first Boreal transgression beds on the lower Ob.

181

33 - 47 ka interval. And finally, 3 out of 6 I4C datings fell outside the indicated range on the section at Karginsky type locality at Karginsky Cape on the Yenisey, where sands with Natica clausa and Macoma baltica are found to separate two tills, i.e. Sanchugovka and Zyryanka moraines. Two of the extreme datings were made on not very reliable materials such as peat (46,000 f 900, GIN-370a) and shells (42,200 k 1000, GIN 387) and only one was made on wood (41,850 +- 1300, GIN 373a). However, Kind (1974, p. 41) wrote: “Another piece of wood from the same trunk but better chemically treated was dated as more than 51,000 yr (GIN 373e)”. The age of plant detritus from the base of Karginski beds fell outside the range, i.e. being older than 50 ka (GIN 369). Nevertheless, Karginski transgression is firmly believed to be Intra-Wiirmian. Marine Karginski deposits seem to be best studied in the Leningradskaya river basin (the river flowing to Toll Bay) in northernmost Asia. Near Schmidt Cape (76”N) Gudina et al. (1983) described a foram complex with Islandiella islandica (Now .) Cassidullina subacuta Gud; Retroelphidium atlanticum (Gud.), Astrononion gallowayi Loebl. et Tapp. from the top of the section in a 50 m terrace not overlapped by a moraine. A warm-water foram complex with the elements of Boreal - Luzitan fauna (Cibicides rotundatus Stsched., Trifarina angulosa (Will.), Elphidiella tumida Gud. ( = E. hannai Cuschm.), Buccella troitzkyi Gud. was also found. A similar complex was found in 40 - 80 m high terraces of Oktyabrskaya Revolutsia Island (79”N); it also includes Boreal - Luzitan species B. troitzkyi, E. excavatum, Asterigerinella pulchella (Phleg.) and others. Makeyev found there a Boreal mollusc Chlamys islandica and the bones of a whale dated b y 14C as 43 ka (Arslanov et al., 1980). As shown by Gudina et al. (1983) and Levchuk (1982), the discussed series of 40- 80 meter high terraces overlapped by Sartan Moraine only in the vicinity on the Byrranga Mountains is” . . , the most thermophilic Pleistocene foram complex, . . . the only known example of the invasion of thermophilic associations eastward along the Eurasian Shelf” (Gudina et al., 1983, p. 95). At present such complexes inhabit the North Atlantic, the areas influenced by the Gulf Stream waters. Findings of Boreal molluscs on Severnaya Zemlya and on Novaya Sibir Island would indicate the penetration of Gulf Stream waters some 2000- 2500 km further to the north in the Arctic Ocean than at present. On the basis of the whole set of data available, the bottom water temperature to the north of Taimyr is estimated to be 5°C and its salinity 35% (Levchuk, 1982). Many researchers drew similar conclusions on extremely large warming in Karginsky time on the basis of spore-pollen spectra from the sections in level I1 and 111 North Siberian terraces (The Quaternary Geochronology, 1980, pp. 183 - 190, 191 - 197, 203, 204, 21 3 - 222, 223 - 229). Thus, according to Saks the time of the second Boreal transgression was the interglacial, being warmest in the Arctic during the whole of the Pleistocene. This conclusion, however, does not agree with the radiocarbon datings of the age of the Karginsky transgression, determined as 50- 22 ka (Kind, 1974; Andreyeva et al., 1981, Gudina et al., 1983). If the age is correct, the question then is why it is that the interglacial occurred only in the Arctic and sub-Arctic, mainly in their asiatic areas at that, while in other regions of the Northern Hemisphere the climate of that

I82

period is known to be harsher than the present climate. It is possible to explain such a paradoxial situation from the point of view of climatology. Furthermore, recently reliable data were obtained for the Norwegian Sea basin indicating that after the 5th isotopic stage, and even after sub-stage 5e, 115 ka ago, the northern Norwegian Sea was constantly ice-covered even during the isotopic stage 3 (Belanger, 1982; Kellog, 1980; Streeter et al., 1982). There is direct evidence that the Arctic climate was not interglacial at all in the period of 30-22 ka. (Vasilchuk et al., 1984). Thus, climatologists and geologists prior t o advancing hypotheses to explain the Karginsky paradox should better investigate the validity of serial radiocarbon datings, giving evidence to the interglacial environment within the Wiirm. Unfortunately there are only a few independent datings of Karginsky deposits by different methods. On the Kola Peninsula, three sections were dated by I4C and Useries methods yielding Ponoi, Middle Wiirm foram complex with Cibicides and Trifarina angulosa (Gudina and Khoreva, 1984). The data obtained (see Table 5.9) were similar t o U-series - D/L analysis results for dating of the J. islandica zone in the sections from Baffin Island, i.e. the 14C age is 28,200- 46,950 years, while the U-series age is 68 - 80 ka (Andrews et al., 1983) and the U-series determined age of “interstadial” Hochstetter in Greenland is 70 - 115 ka. This is clear evidence indicating that the Ponoi transgression and J. islundicu foram zone should correlate with the isotopic stage 5 and not with stage 3d. Thus the author concludes that only a few dozen out-of-range I4C datings of Karginsky deposits are valid out of about 300 and that the hypothesis of the Karginsky thermochron as the Intra-Wurm interglacial is not sufficiently confirmed geochronometrically. In fact, there is a Karginsky interglacial, but it is equivalent t o the Fjesanger interglacial which was Th/U dated as 120 ka (Andersen et al., 1983; Miller et al., 1983). However the upper lacustrine portion of the Karginsky deposits seems to be younger than Zyryanka deposits. The foregoing casts doubt on the problem of the Zyryanka glaciation. The work of Kind (1974) and Troitsky at Karginsky Cape and Arkhipov et al. (1977) in the lower reaches of the Ob showed that the maximum of the Last Ice Age in the Arctic

Table 5.9. Comparison of the numerical age (in ka) of the Boreal transgression by I4C and U-series (after Arslanov et al., 1981) Sampling location, species

‘4c

U-series

Malaya Kachkovka, Kola Peninsula, Cyprina islandica

> 43.7 (JIY - 1360)

114 k 4, (LU 452B)

Svyatoi Nos Bay, Kola Peninsula, Astarte borealis

45.54 t 1.17 (JIY- 137B)

97 4, (LU 455B)

Chapoma River, Kola Peninsula, Cyprina islandica

34.50 t 45 (TA-270)

86 k 3.9, (LU 464B)

*

occurred in the Late Wurm, from 22 to 10 ka BP; this inference is generally accepted. The center of the glaciation was situated on the Barents Sea and Kara Sea shelves, where an ice cap 2.1 million km2 in area, was formed. This ice cap was 3.5 km thick and 4 million km3 in volume (Grosswald, 1983; Grosswald et al., 1978). On the basis of 14C datings of the wood from the beds “underlying a moraine” near Markhida village (the lower Pechora) as 9900 t 100 yr (LU-391) as well as on the datings of tree-trunks from the “till” proper as 9100 -t 100 and 9300 t 60 yr. Lavrov (Arslanov et al., 1981) and Grosswald et al. (1978) claim that this glacier started its advance as early as the Holocene. It is, however, quite possible that Markhida evidence, which is a “compelling proof” according to Grosswald et al. (1978) of the hypothesis on the late culmination of the Wurm glaciation, was incorrectly interpreted. The so-called marginal moraine line at Markhida village can be a pseudomoraine, formed in the beginning of the Holocene when a large block of “dead ice” rapidly melted. This “dead ice” could have been not only of the Late Wurm but also of the Early Wurm. For instance, a pseudomoraine at Denezhkino village at the Yenisey, yielding tree-trunks with ages of from 2425 yr to recent, now moving from the hill with the buried glacier ice of Zyryanka stage of different ages, has the same origin (Astakhov, 1984). As Lavrov (1983) pointed out, the pattern of “Karginsky datings” is characterized by a frequent inversion, when the dates for terrace deposits overlying terraces (20-49 ka and more) are older than those of the till underlying deposits - 23 - 46 ka (Arslanov et al., 1983, p. 33). Lavrov and his co-authors together with Arkhipov et al. (1977) take into account only young under-till datings. This is difficult to agree with, for in such a case I4C datings appear to be adjusted to the hypothesis of glaciation culmination in the Late Wurm. The author (Zubakov, 1972b) published the results of ‘%-dating of fossil wood (15- 16 ka) from the upper part of the sections of young terraces of Yamal and North Siberian lowland. This age contradicts the hypothesis of the Late Wurm glaciation culmination in northern Siberia. Recently Avdalovich and Bidzhiyev (1984) recognized Karginsky terraces topped by peat bogs on the Gydan and Taz peninsulas, which were I4C-dated as 32.7 - 40.7 ka, they pointed out that the terraces were not overlapped by the till. Vasilchuk et al. (1984) carried out detailed studies of a section of the 20-30 meter terrace at the west shore of the Ob Gulf and on the Yavay Peninsula (72”N) and found that syngenetic polygonal-veined ice in its section was formed in the period from 30,100 to 22,700 yr on the evidence of nine 14C datings, which included datings of organic matter from the ice veins. This terrace was not overlapped by till either. Makeyev did not find moraines on the Karginsky terraces with 14C-dated ages ranging from 43 to 19 ka. Moreover, it appeared, that mammoths inhabited the Severnaya Zemlya Islands from 24,960 f 210 (LU 7496) to 19,270 f 130 yr (LU 654b) and from 11,500 +_ 60 to 9950 4 100 yr. The Late Wurm glaciation on Severnaya Zemlya developed in the period from 19 to 12 ka. This glaciation was of the ice cap type, as it is at present (Arslanov et al., 1983). Thus, the culmination of the Wurm Glaciation in northern Siberia indeed occurred after the Karginsky transgression, though it was in the Early Wurm. Two stages of the glaciation can be distinguished. The first stage - the Yermakovo stage -

184

was more than 50 ka BP, as established by I4C dating. It seems to be synchronous with the 4th isotopic stage. The second advance of the Nyapan Glacier on the Yenisey, Lokhpodgort on the Ob, Sartan ( = Zhigansk?) on the Lena was 44- 32 ka BP, from l4C datings of the deposits underlying the till and had several oscillations. The third Norilsk glacier advance, also with a few oscillations, was found to occur from 20 to 15 ka BP. These kryomers alternated with interstadial warmings, though their climatic environment was never better than it is at present. Similar data were obtained for other Arctic regions like northern Canada and Spitsbergen (Lindner et al., 1983; Boulton et al., 1985). The warmest climate was between 60 and 45 ka BP, i.e. the Igarka interstadial. It is difficult to conclude whether a marine glacial eustatic ingression occurred on the northern Siberian lowland (“Boyarka”) and in the lower reaches of the Ob, as it is described by Andreeva et al. (1981) and Arkhipov et al. (1977). Frozen dead mammoths are the characteristic feature of the end of the Igarka thermomer. Mammoth bodies form the so-called “closed system” for 14C datings, thus providing the most valid chronology. Most bodies of mammoths and other large animals are found to belong to the period of 38 - 39 to more 53 ka (Arslanov, Vereshchagin et al., 1980). The mammoth from Lesnaya Rossokha on the Khatanga river (more than 53 ka) was found in the sands overlapping marine Karginsky beds. This excludes the Karginsky transgression from the isotopic stage 3d. Plant remains from the stomach of the Kirgilyakh mammoth (39,570 k 870 yr) give evidence that the animal inhabited the area during a cooling stage. The Nygaard interstadial in Scandinavia, dated by U-series as 52 k 12 ka and by amino-acid racemization as 50-70 ka (Lundqvist, 1983) and the Peter-Bagg more than 45 ka ago from 14C datings in Greenland (Hjort, 1981) can be regarded as equivalents of the Igarka thermochron. The climate during the second interstadial between the Nyapan and Norilsk ice advances (Novonazimovo, Se-Yakha) was cold. This is supported by the first results of oxygen isotope analysis of vein ice from a 20- 30 meters high terrace of Yamal and Gydan. 6 l 8 0 variations by nine samples of vein ice, 31 -22 ka old, comprise 21.4 - 24.8%0,while in the recent vein ice this value is reduced to only 18.3 - 18.7O/uo. This would suggest that the mean annual temperatures were 4 - 9°C lower than at present in the Karginsky interstadial, as Vasilchuk and others (1984) called it, though it would be better to call it the Seyakha interstadial from its type locality. The Tab-Yakha thermomer with an age of 16.5 - 15 ka (Zubakov, 1972b) had a climate similar t o the present one, though with more pronounced continental features. This is indicated by a more northerly shift of shrubby tundra than at present. The Putoran Ice Sheet completely retreated outside the lowland.

Resume (1) The sequence of Pleistocene climatic events of the glacial region is more difficult to reconstruct than of other regions, and that is why it cannot be regarded as an inter-regional standard. (2) The conventional four-unit Pleistocene scheme for the glacial region has been

I85

revised recently. A new pattern of the ice sheet development has been found to be much more complicated than the earlier one, and it seems to have a lot of ambiguities. (3) The glacial eustatic origin of sea-level fluctuations is found to be characteristic of the Pleistocene World Ocean, the Arctic Ocean included. No separate transgression occurred in the Arctic Ocean during the Pleistocene. (4) The hypothesis of the Pan-Arctic ice shelf seems to be not confirmed by stratigraphic-chronological evidence, at least as it concerns the Late Wiirm, though it is on the Late Wiirm evidence that this hypothesis is based.


SECTION 111

THE HISTORY OF CLIMATE THROUGH THE PLEISTOCENE


Chapter 6 ON THE TIMING OF PALAEOCLIMATES IN THE PLEISTOCENE

6.1. Debatable problems of inter-regional climatostratigraphic correlation 6.1.1. On correspondence between the numbers of climathems on land and in the sea

The comparison of the regional climatostratigraphic units (Table 6.1) has shown that they contain an equal number of climatic events. In particular, the number of climathems in the most complete schemes of the continental Pleistocene is almost the same as in the oxygen isotope scheme of the deep-sea Pleistocene. It has been revealed that even the most weakly pronounced isotope stages have their analogues in the continental schemes. Thus, oxygen isotope stage 14 corresponds to the Oka kryomer of the Russian Plain and the Wilga kryomer in Poland. Isotope stage 17 has its analogues in the loess sequences in Czechoslovakia, in the Sea of Azov region, the Don Basin, Soviet central Asia, China as well as in southern Poland and East Germany. Stage 20 has been correlated with the Pokrovka- Helme kryomer (glacial A) in Europe. The Linge - Elmskhorn kryomer is comparable with isotope stage 24. The correlation allows us to reject as logically inconsistent the statement of some researchers (Nikolayev and Nikolayev, 1984; Blyum, 1982) that the number of known warm stages is smaller than that of isotopically light stages in Shackleton’s scale and that some of these isotope stages (for instance, stage 7 ) reflect the process of freshening of surface oceanic waters due to surge effects. The analysis has shown that the succession of oxygen isotope stages adequately reproduces the sequence of global climatic changes and it can therefore be used as a standard. 6.1.2. On two stratific lines in the geo-historical classification of the Pleistocene

The modern state of chronology and classification of the Pleistocene is full of contradictions. On the one hand, the bulk of information on the local, exceptional features of the Pleistocene history has been rapidly increasing and there have emerged a legion of new local stratigraphic names. On the other hand, the absence of a general stratigraphic scale for the entire world or even for the USSR territory, without which palaeoclimatic reconstructions are impossible, has been more and more acutely felt. Meanwhile, as has been seen in the previous Section 11, the stockpile of information has already made it possible to outline the contours of a global geochronological Pleistocene scale, which can be highly accurate and detailed and at the same time quite different from the traditional Phanerozoic scales. The practice imperatively requires a fast solution to this problem. However, a unified strategy has not yet been developed in this field and a solution to this problem depends now on two approaches and two philosophies validating them.

Table 6. I . Inter-rogional climatostraligraphic

COI r e l a h i

of the Lower and Middle I’lciitocene 0

Loess area

~

-

Se

_

_

Ukraine - Don

_

_

Priluki

Eltigen

~

_

Glacial

Czech. Hungary

_

Mezin s.

BI

MB

5

Geroevskoe 11

y"

2

ZavetninoGeroevskoe I ?

Girkan

C

P,,

Tyasmin - Merzalov I .

~

80 m terr.

I

i ' .'I Nomentanan

Babel

=I I

k

(Mindel 111)

I20 m. terr.

1

Singil fauna Lower Kharar

=

2

2

'Ozhki

Kamensk s.

e 3s

Lubny - lnzhavino s. Sula

~

Sefa

Martonosha

-

-

$

Port Katon

,"

Likhvin s.

B

O k a - Dainava

m

D o n - Pereksha - Krukenichi

O"

S6

H

LS

z

l6

2

- 2 S6'

1

c

c

Il'inka - Sokal s

.

Apsheron Middle Apsheron

e

3

.

SIC.

Roslavl' - Shklov - Muchkap

1

2%

7

I

E .

-

Kana

Novokhoperrk

16'

SetUn' - Lipetsk

Sl

Akulovo - Korchevo

17

Lika - Narev?

I

Pokrovka I.

Regrescion

Upper

Chumbur

,

s5

- 5I5

a

~j

Koshnitsa

Margaritovka

-.-s

Chekalin - G r o d n o

5

C

Vorona s.

Priazov loess

Lower Baku

Karai - Dubina

12

5

PD2

Balashob soil

3

Chumbur I. Urzuf so11 Il'ichevsk (?) I.

5,

,,, J

PD,

I L 6

DV

=

f I

1 '$

-6

E"

c

--

F

Rorostelev I.

x

30 9k

!--

PDI

Miuss

2

Q 0

E

~

m

Cawan

-

Tiligul - Borisoglebsk I .

I E

Upper Baku

-~

.Y

Ficarairi

Zavadovka

__

Taganrog

m

Kalinin? ? Odintsovo - Duna

Dn ie p er

-?

?. o

m

. - c

Romny s.

Elanchik

I

3

-

53

b n

Urundzhik

Patrai

16

Flarninian erosional phase ( = Mindel I ? )

. DB

Shapsugo

~

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192

The first approach formulated by Hedberg and expressed in the Stratigraphic Guide (1976) can be defined as a chronostratigraphic one. Its principal feature is the selection of stratotypes for the stratigraphic boundaries, which in the long run is accepted “by agreement”. Following recommendation of the Guide, all regional stratigraphic units should be associated in any possible way with the boundary stratotypes, the so-called “gold spikes”. An example showing successful application of this pragmatic approach was considered to be the identification of a new Pliocene - Pleistocene boundary by the stratotype of the Vrica section. The second approach proceeds from the experience of the European stratigraphic researchers. In can be found in the works by Librovich (1954), Menner (1962), Schindevolf (1970), Sokolov (1978), Krasilov (1977, 1985) and others, and also in the documents of the USSR Stratigraphic Committee. This approach is chiefly based on the search for natural geological historical boundaries as they are globally fixed and traced through ecological reconstructions. For short, it can be called a signal (Zubakov, 1978b) or event (Valliser, 1984) stratigraphy. According to this approach, in order to find valid boundaries of the global stratigraphic units it is necessary to have not only one “gold spike”, which can easily be driven into the wrong spot, but the greatest possible number of standard sections (reference points) in various facies zones, different basins and at different latitudes, which should be reliably synchronized. Following this approach, the identification of the Pliocene - Pleistocene boundary with the stratotype in the Vrica section is an example of a stratigraphic error. The lower boundary of the Pleistocene as identified by Lyell (1840) was conventionally assumed to coincide with the base of the Sicilian regional stage (Gignoux, 1950) and in the USSR with the base of the Baku and Chauda regional stages of Paratetis (Andrusov, 1965; Markov et al., 1965; Fyedorov, 1957, 1963). As was suggested by Zeuner at the International Geological Congress in London in 1948, the Pliocene - Pleistocene boundary should be found in the lower layers under the base of the Calabrian and Villafranchian. However, in 1952 Zeuner (1959) admitted this statement to be erroneous. Subsequently, Selli (1975) showed that the base of the Villafranchian is older than that of the Galabrian. Still later on, Ruggieri and Sprovieri (1977) proved that the Calabrian layers are identical to the Sicilian ones. Thus, it turned out that IGC recommendation (adopted in 1948 and revised in 1972) proceeded from wrong premises. However, the idea about shifting the Pliocene - Pleistocene boundary, to be more accurate the Neogene -Quaternary boundary, to a lower position, proved to be very tenacious because of an obvious discrepancy between the Quaternary system and other Phanerozoic systems. However, in 1984 a new recommendation has been adopted, which advised that the sapropel layer “e” in the marine section of Vrica, Southern Calabrian, with calculated age of 1.64 Ma (Aguirre and Pasini, 1985), should be taken as a stratotype of the Pliocene - Pleistocene boundary. This solution was in complete conformity with Hedberg’s chronostratigraphic concept, which rejected palaeoclimatic data as instrumental for geohistorical classification. This fact might be considered very distressing since it is next to impossible to correlate the Pleistocene continental sections by means of the “gold spike” of Vrica. Because the Vrica section has not been climatostratigraphically studied at all, the only instrument for such a correlation is palaeomagnetic data.

193

However, the interpretation of magnetostratigraphic data for the Vrica section is very ambiguous. By interpolating the dated micropalaeontological levels such as LAD Discoaster, FAD Gephyrocapsu oceanica and others, the authors of the recommendation have identified the polarity zone lying some 7 to 10 m below the sapropel layer “e” with the Olduvai aged at 1.67 - 1.86 Ma (Colalongo et al., 1982; Rio, 1982). But the comparison of the Vrica section with other sections of the marine Pliocene in Italy and in particular with the sections in Santerno (Arrias et al., 1980) gave rise to the conclusion that the indicated normal polarity zone can correspond to the Reunion one, which is 1.98 - 2.13 Ma of age or even to an older zone dated at 2.33 Ma. The chronometric dates available on the Vrica section are in a broad range of 1.99 k 0.8 Ma and do not contradict the results of the above comparison. The interpretation of the dating levels in the Vrica section is also ambiguous. Thus, FAD Gephyrocupsu oceanica found 25 m above the sapropel layer “e” is estimated in the deep-sea sections either at 1.3 Ma (Rio, 1982) or at 1.77 Ma (Berggren et al., 1980). Some scientists date FAD (Hyalinea bultica at 2.2 - 2.5 Ma both in the Mediterranean (Zagwijn, 1974) and in the Pacific and Indian Oceans (Van Goersel and Troelstra, 1981). And the presence of Cytheropteron tesrudo in the Vrica section at a level of 1.6 Ma is by no means its first appearance in the Mediterranean, as has been shown in Ruggieri’s later investigations (Rio, 1986). Thus, the position of the Vrica “gold spike” proved to be very insecure and the uncertainty range from 1.64 to 2.5 Ma is too great. This is the first point. Secondly, and this is most important, let us remember that the stratigraphy of the continental Pleistocene (as well as of the Upper Pliocene) is chiefly based on palaeoclimatic data, which are indispensable for accurate and reliable land - sea correlation. The climatostratigraphic position of the Pliocene - Pleistocene boundary under the Sicilian - Chaude - Baku stages and the tills of the first continental glaciation in Europe is completely justified. It should be explained that in the USSR the first, reconnoitering in essence, palaeomagnetic measurements (Gromov et al., 1969; Menner et al., 1972; Zubakov and Kochegura, 1971; Pevzner and Chichagov, 1973) have rise t o an erroneous idea that the base of the Baku and Chauda sediments is younger that the Bruhnes-Matuayma boundary and does not coincide with the boundary under the Sicilian - Menap - Nebraska adopted in Europe and America. The latest studies have shown that the age of the base of the Chauda regional stage in the Georgian parastratotype sections is 1.1 Ma (Zubakov et al., 1975). In the Sea of Azov region the Chauda lower layers are correlated with the finds from Tamanian assemblage of mammals (Lebedeva, 1972), which independently confirms the Matuayma age of the Chauda base. In the Azerbaijan sections, the Tyurkyan formation, where the N/R reversal is presumably of the Brunhes - Matuyama age, has recently been dated by ash through the fission-track method at 0.95 - 1.050 Ma (Ganzei, 1984). And finally, in 1985 Mamedov and Aleskerov (1985) found in drill cores from the Tyurkyan site in the Kura Valley Diducna purvulu that dominates in the Baku stage (Geochronology . . ., 1985), Thus, the latest data have confirmed the earliest conclusion of Popov and Rodzyanko (1947) and others that the Chauda and Baku sediments are synchronous. But the age of the base of the Chauda-Baku stage turned out to be 1.0- 1.1 Ma, i.e. 0.4 Ma older than it was supposed earlier.

194

At that time it was also found that the age of the base of the Sicilian stage is 1.15 Ma (Colalongo et al., 1982; Ruggieri et al., 1984). According to recent data, the Menap glaciation took place before the Jaramillo event between 1.2 and 1.1 Ma (Zagwijn and Doppert, 1978; Zagwijn, 1985). The age of the Nebraska B tills of North America (Easterbrook and Boellstolfe, 1981) and of the tills left by the greatest glaciation in the Patagonian Andes (Mercer, 1976) has been estimated at about 1.2- 1.0 Ma. The last great reorganization of the organic world, i.e. the replacement of the Villafranchian mammalian assemblage by the Tiraspolian - Galerian one, that is the proper Pleistocene fauna, also occurred between 1.3 and 0.9 Ma (Zubakov, 1974; Azzarolli, 1983). The genus Homo also emerged in this time period (Ivanova, 1965). Finally, the erosional unconformity in the Alps assumed by Penk (Fink, 1974, 1975) for the beginning of the Diluvium - Pleistocene and similar unconformities in the mountains of Soviet central Asia (Dodonov and Ranov, 1984) and the Altai (Borisov, 1984) have been dated palaeomagnetically at between 1.1 and 0.9 Ma. We can only wonder how - long before accurate dating and correlation methods appeared - our predecessors could find a synchronous stratigraphic level in different areas and unanimously identify it as the Pliocene - Pleistocene boundary. Thus, it is the history itselfthat has confirmed the efficiency of methodologicalprinciples of the event stratigraphy. We might only add that this boundary is also perfectly identifiable in the deep-sea section by the appearance of small Gephrocapsa dated at 1.13 Ma (Rio, 1982), Mesocena elliptica dated at between 1.3 and 1.0 Ma (Berggren et al., 1980) and M . quandrungula dated at 1.1 to 1.8 Ma (Bukry, 1982). All these considerations and the information cited allow us to make two conclusions: (1) The conventional Pliocene - Pleistocene boundary based on climatostratigraphic principles turned out in practice to be more significant, and more realistic and convenient than the boundary adopted “by agreement”; (2) The strategy itself of determining stratigraphic boundaries by some signal or event, and in particular by temperature trends, appeared to be more vital and practical than the recommendations of the American theorists of stratigraphy. It is interesting to mention that a similar situation arose when the stratotype of the Pleistocene - Holocene boundary should have been agreed upon. One of the three sections in southern Sweden, namely in Gothenburg, Solberga or Moltemyr should have been selected for this purpose. However, the members of the Working Group could not come to an agreement either in selecting the stratotype, or in determining the strategy for the development of the global Holocene scale (Olausson, 1982, Konigsson, 1984). 6.I .3. Comparison of experiences in the long-distance stratigraphic correlation of the Pleistocene During the last 20 years, three attempts have been made at inter-regional correlation of the Pleistocene for the USSR and some foreign territories, which have been described in three collective monographs, namely in three volumes of The Quaternary Period by Markov, Velichko, Lazukov and Nikolayev (1965, 1967), in the third

195

volume of Geochronology of the USSR edited by Zubakov (1974) and in two volumes of Stratigraphy of the USSR. The Quaternary System (1982- 1984) edited by Krasnov (1984) and Shantser (1982). The principles of correlation, i.e. the strategy, adopted in these studies are different and it would be useful to discuss them. The first work mentioned above presents in every detail the palaeographic concept of stratigraphic correlation, which is now adopted by many scientists (Lazukov, 1980; Svitoch et al., 1978, 1980; Veklich, 1982; Voskresensky et al., 1984). The most important place is here given to the problem of synchroneity and metachroneity of climatic events in different geographical environments. As long ago as 1938 Gerasimov and Markov (1938) set forth a hypothesis about metachroneity of glaciations, which was further developed by Markov and Velichko (1967). The concept of metachroneity was directed against “metaphysical principles of synchroneity” . According t o the authors of this concept, it shows the logical ways of revealing the true paIaeogeographic principles of inter-regional correlation of glacial (or, in general, climatic) Pleistocene events in the areas, where the types of atmospheric circulation differ radically from each other (Markov and Velichko, 1967, pp. 83 - 84). The authors start with the analysis of uniformitarian principles and give an example of the present differences in the glaciation of northern and southern Alaska. In the north of Alaska, precipitation makes up from about 100 to 200 mm a year and the annual temperature is about - 14°C (Point Barrow), whereas in the south precipitation is between 2,000 and 3,000 mm and temperature +S.7”C (Sitka). With a reference t o Vaskovsky’s data ( I 959) on mountain glaciers in north-eastern Asia, Markov states that during ice ages in mountainous areas at the level of ancient glacial drift the July temperature was 2” to 2.S”C lower than at present, whereas winter was milder and the mean annual temperature was 2°C higher than the present one. The amount of precipitation increased because of heavier winter precipitation. On the basis of such examples, the conclusion was drawn that in order to make a correct global correlation it is necessary “. . . to compare the glaciations (of northeastern Asia) with more profound marine palaeogeographical environment than the present climate of north-eastern Asia” (Markov and Velichko, 1965, p. 94), because . . . “the growth of glaciers was promoted by milder winters, less intense winter anticyclone and heavier precipitation, i.e. the interglacial environment of the opposite Atlantic coastal region” (ibid, p. 95). Markov’s general conclusion was the following: “. . . 1) A colder climate of north-eastern Asia promoted the development of underground ice and impeded the development of surface glaciers; 2) climatic cooling and the development of the surface glaciers in north-eastern Asia were not synchronous; 3) the development of surface glaciers in north-eastern and north-western Eurasia was also asynchronous” (ibid, p. 95). Although the geographical points of this conception cannot raise any objections, the factual information gathered in this book has not confirmed it. On the contrary, the first radiological dating of tills and interbedded layers in the mountains of northeastern Asia and Alaska clearly shows that mountain glaciers over the Kolyma Plateau (Voskresensky et al., 1984), the northern coast of the Sea of Okhotsk (Ananiev et al., 1985) and the Seward Peninsula (Hopkins, 1973) developed synchronously with North Atlantic glaciers. Tomirdiaro (1985) and Kolpakov (Biske,

196

1984) indicate that the till and fluvioglacial complexes of ice piedmonts are replaced

in facies by Edoma formation with underground polygonal-vein ice and not alas by interglacial deposits as should be expected according to the concept of metachroneity. Thermoluminescence dates (Table 5.8) show that dark coniferous taiga with some broad-leaved species developed in north-eastern Asia during global climate warmings. These alternated with mountain glaciations, when the climate was colder than at present. Thus, we cannot share the statement of Markov and Velichko (1967, p. 91) that “. . . glaciation follows the course of precipitation and is reverse to the temperature course”. On the contrary, the information presented in Section I11 convincingly shows that both surface and subsurface glaciations develop synchronously with variations in temperature. At the same time it has been revealed that the type of glaciation is determined by atmospheric precipitation and, therefore, it differs from place to place. Both continental ice sheets and mountain glaciers developed asynchronously and their greatest advances occurred during different periods in different geographical regions (Zubakov, 1963). All this was perfectly demonstrated by the American scientists, who studied the Late Wisconsin ice sheet (see Section 5 . 5 ) . Consequently, in the Pleistocene the dynamics of glaciation in each region of the glacial zone had its specific features. This means that the development of glaciations is characterized not by metachroneity, but by spatial differences caused by palaeoclimatic factors pertaining to its nature and dynamics. This, in particular, allows us t o identify different spatial-dynamic types in the glacial development (Zubakov, 1965). Thus, the advantage of palaeogeography as compared with stratigraphy cannot be justified methodologically. Palaeogeography tends to simplify the history of the Pleistocene, depicting it as a priori understandable. It gives a minimum number of glacials and interglacials, which contradicts the present data. Therefore, the monograph under consideration, which was undoubtedly the greatest step in the Pleistocene studies, today cannot be an example for developing a new system for the Pleistocene classification. It is not easy to understand the principles for the global stratigraphic correlation of the Pleistocene events given in two volumes of The Quaternary System (Shantser, 1982; Krasnov, 1984). These principles have been developed as a combination of three elements: the well-known experience of regional stratigraphic meetings, Krasnov’s theoretical ideas and Nikiforova’s views, which in many respects differ from each other. As is known, the stratigraphic schemes worked out by regional meetings are either based on the collective scientific opinion that is dominant at the present time or are often the result of an agreement by mutual concessions. These schemes are constructed by generalizing information step-by-step from particular to general. In the above-mentioned work Krasnov compares these schemes without any serious analysis or generalization (Krasnov, 1984). The theoretical ideas of Krasnov are mostly notable for the taxonomic aspects of stratigraphy that are set in the first place. This taxonomy is based on the duration of events, which, after Krasnov (1973, 1974), can be most clearly seen in the solar radiation curve. Krasnov, together with Shantser (1982), suggests five units for the general stratigraphic Pleistocene scheme. These units have fixed (except for the

197

boundaries) duration: a zveno of 200 ka to 300 ka, a step ( = climatolith) of 20 to 100 ka, a stadia1 of 5 to 10 ka and a level of 1 to 5 ka. It is suggested that zvenos are formed by great glacials and interglacials, steps by small glacials and interglacials, stadials by great stages and interstadials, and levels, by small stages and interstadials. At the same time the authors of The Quaternary System (1984) d o not even touch on the question of how the global steps and levels with isochronal boundaries should be identified, if their volume is determined by the local sub-units with the boundaries that are transgressive in time. It can be understood from the chapter written by Nikiforova (in Krasnov, 1984) that these units should be identified through the stratotypes of the steps or the global climatostratigraphic horizons. It is also stated that all the horizons (!), i.e. the cold and warm steps of extraglacial zone, have faunistic characteristics. It is indeed that in Table I1 (Appendix to the first half of Volume 2, Krasnov, 1984) all the horizons up to the Kolkotovaya, counting from bottom to top, are characterized by the dominant species of freshwater molluscs and rodents. Such a super-accurate diagnostics of the steps that lasted from 20 to 100 ka gives rise to great doubts. It looks more like a formal assignment of the species to certain steps, which actually occurred within a much greater time interval. The two-volume edition of The Quaternary System (1982- 1984) is valuable first of all because of the factual regional information presented by a large body of specialists in local geology. As to the sections devoted to inter-regional correlation and classification of the Pleistocene, we are sorry to state that they do not give any clear indications for geohistorical classification. A combination of theoretical ideas of Krasnov, who on the whole follows the climatostratigraphic concept, and Nikiforova, who adheres to the recommendations of the chronostratigraphic school, forms a contradictory view that cannot lay grounds for a methodologicallyfounded strategy.

6.2. Rhythm-chronological approach to the Pleistocene classification

In The Geochronology of the USSR (Zubakov, 1974) and in the works by Zubakov (1961, 1963, 1973, 1968b,c; 1978a,b) an approach to the Pleistocene classification has been developed, which can be described as the rhythmchronological approach. It is based on the following principles: (1) In the climatic-sedimentary cycles, the stratigraphic boundaries that are almost isochronal can be identified, which follow changes in the global temperature trend. These boundaries are not so distinct as the boundaries of the local climaticsedimentary units, which are usually clearly pronounced but transgressive in time, although in some cases they can also coincide with the latter boundaries; (2) These boundaries can be observed and recorded only with the help of a complex methodology, including lithological, palaeontological, geochemical, isotopic, palaeomagnetic and chronometric techniques. In this case only “terrestrial” dating can be made, not the correlation with an insolation curve; (3) Such boundaries cannot be established by agreement, having an example of

198

some individual section (like the “e” level of Vrica). They are identified by selecting climatostratigraphic signals that are observed and dated in the continental, shelf and deep-sea sediments; (4) The general climatostratigraphic scale based on these principles will be chronological and “flexible”. The valid marks will stick up without any spikes, their unique properties and first of all numerical age being constantly confirmed and made more precise. 6.2.1. On three types of time classification of the Pleistocene climatic events

Although climatostratigraphic units of the Pleistocene were geologically synchronous, their local nomenclatures are hardly comparable for many reasons. First of all, there are gaps in the geological record and different authors give subjective interpretations of the climatic events that occurred in different regions. However, of great importance are also objective differences in the climatic history of the Pleistocene in low and high latitudes, in the ocean, the shelf seas and on land. Because of these reasons, there appeared three versions of timing the Pleistocene climatic events, namely the glacial, the Mediterranean and the deep-sea systems (Table 6.2). These systems differ in the scope of time and the number of units. Thus, the classical glacial system based on the study of glacial assemblages in the middle latitudes of Europe and North America within the conventional Pleistocene boundaries determined at the base of the first continental ice sheet in Europe distinguishes six glacials and five interglacials (Yakovlev, 1956; Flint, 1957; Moskvitin, 1967; Woldstedt, 1958). The duration of glacials exceeds by far that of interglacials, the former usually estimated as 10 to 15 ka, but no more than 30 ka (Liittig, 1965b; Moskvitin, 1967). The traditional Mediterranean system (which is also applicable to the shelf seas) recognizes three or four interglacial transgressions separated by two or six erosional regressive phases (Zeuner, 1959, 1965; Ambrosetti et al., 1972). However, this version suggests that thermomers should be longer than kryomers (Table 6.2). Thus, according to recent data, the Mediterranean Riss - Wiirm, which corresponds to the Strombus horizon, is about 140- 150 ka long (or otherwise, estimated at 220 to 80 - 70 ka), while the duration of the Riss - Wurm in north-western Europe (the Fjmanger - Mikulino thermochron) is estimated at only 10 to 15 ka (Mangerud et al., 1981). The duration of the Riano ( = Palaeotyrrhenian) thermochron is about 170 ka and the corresponding Holstein - Likhvin does not last more than a few thousand years. In the deep-sea Pleistocene classification based on variations in the isotopic composition of oxygen in shells of planktonic foraminifera in the equatorial Pacific (Shackleton and Opdyke, 1973, 1976), the number of climatic events is more than twice as great as in the conventional system with kryomers and thermomers of almost equal length. As has already been shown, all isotopic stages have their analogues both in the glacial and Mediterranean sections. The taxonomic rank of some of them will be estimated further. Thus, it can be seen that the three versions of the Pleistocene classification do not contradict each other; moreover, they even supplement each other. Being tax-

199

onomically different, each of them reflects various aspects of changes in the global climate. The oxygen-isotope curve represents on the whole a general state of the Earth’s glaciation, which can theoretically be asynchronous in the Northern and Southern Hemispheres because of the mechanisms of precession that determines the Table 6.2. Three types of paleoclirnatic Pleistocene periodization

I

Deep-sea

1 Glacial type

stage)

Pontino erosional phase

w I I - 111

2-4

Large Wurm Wisconsin

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%

5e

6

.

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Small R - W lino)

(=

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‘L

7

Ostian erosional phase

Riano\

i Large Riss Illinoian

I

.m

rq r

Tarquinio - Uzunlar

M

PraeRiss Likhvin - Yarmouth

14

Flaminian erosional phase

16- 18

Large Mindef “Kansan”

19

20

Atern

21

Sicilian - Portuensian Chaudian

22

-

I

I

EL

23

m5j

I

2s

Bavel

26

Menap

... 30

Leerdam Linge

24

Cassian erosional phase

Dorst

?

I

Large Gunz “Nebraskan”

200

opposite fluctuations in insolation in both hemispheres. Therefore, it is possible to use the oxygen-isotope curve for determining the length of the global climatic events. However, it does not permit the division of the Pleistocene climatic history into natural stages, i.e. it does not permit t o classify the Pleistocene by the events that have once taken place (to establish the so-called event classification of the Pleistocene). 6.2.2. The role of the 400 ka cycle f o r chronological classifcation of the Pleistocene Any natural system of time classification is more preferable than a devised one. The only way to develop such a global system of timing the Pleistocene is t o correlate globally the climathems and t o identify the greatest steps in the evolution of the natural environment. Therefore, it is necessary to estimate (weigh) palaeontological characteristics of different climathems in order to reveal the most reliable among them. On the other hand, the regular principles of the recurrence of climatic events in time should be determined, i.e. it is necessary to understand whether there is any natural rhythm in the course of global climatic events. Milankovich (1930) recognized three orbital rhythms affecting the redistribution of insolation at different latitudes, namely changes in the obliquity of the ecliptic ( E ) through a range of 21'39' to 24'36' over a period of about 41 ka, in the precession of the equinoxes ( E ) with a period of about 21 ka and eccentricity of the Earth's orbit (2) with a period of 92 ka. It turned out that fluctuations in insolation produced by these cyclic changes (three rhythms) agree satisfactorily with the empirical oxygen-isotope curve. Consequently, the latter is causally related to insolation fluctuations, and the dates of isotope peaks can be correlated by astronomical calculations (Emiliani, 1978; Hays et al., 1976; Imbrie et al., 1973; Komintz et al., 1976; Morley and Hays, 1981). It might be thought a priori that there are climatic rhythms with greater length and amplitudes, which could be valuable for palaeoclimatic classification. Geologists have long noticed the traces of such cycles, but have been uncertain as

J

Fig. 6.1. 370-380 ka climatic rhythm revealed in the succession of glacial and interglacial events (from Zubakov, 1968). The glacial curve in thousands of years. I - XI11 - cyclic marine terraces with heights in meters: I 2-3, I1 - 3-4, III - 5-6, 1V - 15-25, V - 30-35, VI - 40-50, VII - 50-65, VIII - 70-80, IX - 85 - 110, X - 110- 125, Xi - 135 - 150, XI1 - 155 - 170, XI11 - 190-250(after Kaiser, 1965). W , R, M, G, D - symbols of Alpine glaciations; RW, MR, GM, DG, BD - interglacials.

20 1

to their duration. In 1966 Balukhovsky and the author independently advanced information on the traces of a 330 ka cycle (Balukhovsky, 1966 - 1973) and 370 - 380 ka cycle (Zubakov, 1967 - 1968) in the Pleistocene sequences (Fig. 6.1). Subsequently, Sharaf (1974) made calculations similar to Milankovich’s, but covering a time interval which was 30 Ma longer, and identified new astronomical rhythms associated with changes in the tilt cycle having a period of 200 ka and in eccentricity over periods of 425 ka and 1.2 - 1.3 Ma. These data have further been confirmed by A . Berger (1978). The 400 ka cycle has recently become the subject of many publications, which appeared after the works of Briskin and Berggren (1975) and Briskin and Hare11 (1980), who revealed this cycle in the deep-sea cores (Fig. 6.2). According to the author’s hypothesis (Zubakov, 1968a,c), a climatic cycle of 380 ka called “zveno” includes up to eight phases (Fig. 6.1). It has been suggested that every phase is characterized by distinctive climatic features (as the reflection of orbital position) and in sections, by biostratigraphic features. Table 6.3 shows the development of this hypothesis. Climatic events are presented in such a way that four time intervals of 380-420 ka can be compared with each other. Any event, thermomer or kryomer can be selected to begin the count of the rhythm. In his

Fig. 6.2. Graph showing faunal indexes T,,,and T, of estimated temperatures and seasonality (7, - 7,) versus time. Paleomagnetic boundaries and Ericson’s faunal zones are inserted near the time in Ma. Four major roughly symmetrical climatic cycles are revealed in T,,,.The lowest estimated winter temperatures occur at the base of Jaramillo and the warmest estimated winter temperatures in the middle and upper Brunhes. Three points weighted moving average delineate the major temperature pattern. The seasonality shows an inverse trend to T,,, and T, (after Briskin and Berggren, 1975, fig. 10).

202

previous works the author tried different points of counting off. In the suggested version, the last incomplete 400 ka cycle is assumed to begin with the Dnieperian - Saalian ice advance and isotope stage 10 (Table 6.3). It turned out that the last complete cycle contains stages from 20 to 1 1 with relevant continental analogues and the last but one cycle comparises stages from 30 to 21. In the table we have composed, one can notice the following interesting features: ( 1 ) Within all three rhythmic patterns the most extensive glacial eustatic transgressions, the highest salinity of the shelf seas and the greatest invasions of stenohaline and warm-water marine molluscs can be observed at the same level, i.e. within substage 5e and stages 15 and 25. Thus, it was only within stages 5 and 15 that the Mediterranean stenohaline fauna reached the eastern part of the Manych Strait (the Eltigen and Patrai transgressions). In the North Atlantic and the Arctic, stage 5 coincides with the warmest Late Pleistocene transgression (Fjersanger - Ponoi “Kazantsevo - Karginsky”). At that time the Gulf Stream waters spread into the Arctic by two or three kilometres as far as the New Siberian Islands. Stages 15 to 13 are associated with the Ust Solyenaya - Ob layers containing warm-water Miliolinella pyriformis and Glandulina nipponica, whereas stages 25 to 23 correspond to the Anvilian - Enmakai layers with the warm-water Elphidiuirn quasioregonensis and dextral Neogloboquadrina pachyderrna. (2) Marine transgression, which occurred within stage 7 (Salemal- Pupkovo Kresta - Kotzebue), stages 27 and 17 (Portuensio, Tiltim - Bolgokhtokh - Pinakul- Upper Olkha) were much colder and the water salinity of the Mediterranean, the Black Sea and the shelf Arctic seas was relatively low. In the Arctic and the Bering Sea these transgressions were accompanied by the glacial - marine facies and poor assemblages of cold-water molluscs, foraminifera and ostracods. (3) For all three time intervals the greatest peaks of heavy isotope seem to be at the same levels. To be more exact, they are found at two levels (the lower level is taken by stages 6, 16 and 26, and the upper one by stages 2, 12 and 22). The kryomers of the upper level in the Northern Hemisphere are marked by not very great advances of the ice sheets into the middle latitudes. However, the loess sequences at this level (the Wiirm, Tiligul, Trostnyan) reach the greatest thickness. The climate of the Northern Hemisphere was most severe and highly continental within stages 2, 12 and 22. The maximum southward advancement of the ice sheets in the Northern Hemisphere occurred during stages 6, 16 and 26 (the MOSCOW, the Don) and the preceding stages 8 - 10, 18 - 20 and 28 - 30 (the Saale - Dnieper, the Lika - Setun and the Nida - San, the Nebraska A,B) which means that the climate at that time was very “snowy”. (4) It turned out that the evolutionary steps of the development of the organic world are in a certain relationship with the stages of the conjectural 400 ka rhythm. Thus, the last appearance of the dominant forms of coccolithophorids is observed within stages 12 (LAD Pseudoeniiliania lacunosa), 22 (LAD Mesosena elliptica, small Gephyrocapca) and 32 (LAD Helicosphaera sel/i). Stage 22 is marked by a replacement of the Epivillafranchian - Tamanian fauna by the Galerian - Tiraspolian fauna. Within stage 12 the indicated faunal assemblages are replaced by the Svanscombe - Volga fauna. A complete extinction of mammoth fauna took place at the end of stage 2.

Table 6 . 3 . Synphase correlation of local climatostratigraphic subdivisions in four subsequent 400 ka climatosedimentation cycles.

N w 0

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The major events of man’s evolution are also associated with the same phase of the rhythm: for example, Homo sapiens appeared within stage 3, its Acheulian ancestor emerged within stage 12, Homo erectus with its chopped artifacts (Azykh, Vallonnet, Kuldara) developed within stages 22 - 24 and Homo gen. in the Olduvai event, within stage 32, 1.2- 1.3 Ma BP. ( 5 ) A similar picture can to a certain extent be observed in the geomagnetic field within three cycles under consideration. Thus, the Levantine excursion (330 - 295 ka), the Brunhes - Matuayma reversal (734 ka) and the Lower Cobb Mountain reversal (1.1 Ma) are divided by almost equal time intervals of about 400 ka. One might have thought that these major polarity events of the Pleistocene took place at the time when the greatest (by volume and extent) ice sheets were forming on land. Excursions r5 (Jamaica - Biwa I, 210 - 180 ka), r8 (Yakhno - Don, 550 - 600 ka) and presumably the upper Cobb Mountain reversal are also separated from one another by time intervals of 400 ka and took place at the time when the volume of continental ice was increasing. A series of the Late Wiirmian excursions (rl, r2, r3, namely Gothenburg, Mono, Laschamp) dated at 13 - 43 ka, excursion r7 (Ureki - Sneik River, 400 - 470 ka), presumably the Zykh - Taylor Valley excursion (with K/Ar age of 840-830 ka) and the upper Jaramillo reversal (900 ka) are located in almost the same order within kryomers 2 - 3, 12 and 22. The rhythmic analogues appeared to be the Blake and Upper Jaramillo reversals. We have already given enough examples and information to see that there are certain natural stages, or better to say steps, in the succession of climatic events and induced evolutionary changes in the organic world, when the rate and intensity of certain natural processes increased or changed abruptly. Such steps marked by the evolution of the organic world are associated with boundary isotope stages 2/ 1, 12/11, 22/21 and 32/31. More than twenty years ago the author wrote “. . . a zveno cycle represents a kind of a periodic table of climatic fluctuations, since each stratigraphic unit within this cycle is maked by specific features and certain length. Taking this into account, the author thinks that the zveno cycle must be the largest climatostratigraphic subdivision of the international scale, which can be identified within a stage” (Zubakov, 1968c, p. 52). There is some confirmation of the prognostic role of the 400 ka cycle. For instance, the age of the moraines in the Don Basin, which as early as 1968 were recognized as the Dnieperian, has subsequently become 400 ka older. The former Upper Pleistocene interglacial layers in the vicinity of Roslavl, where the Tiraspolian fauna has been revealed (see Section 5.1), have now been identified as Middle Pleistocene. A great many kryomers have also been updated and new tills and glacial horizons have been discovered. Thus, the 380-400 ka cycle actually appears to be the greatest climatostratigraphic subdivision of the Pleistocene. On the other hand, it is the smallest biostratigraphic unit that can be identifed globally. For example, by the Mediterranean - Black Sea regional stratigraphic scale, the last zveno corresponds ’The Stratigraphic Committee of the USSR has introduced the term “zveno” as the greates unit of the Quaternary System into the USSR Stratigraphic Ctassification (1 978).

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to the Tyrrhenian - Karangatian s. lato regional stage, the last zveno but one to the Tarquinian - Uzunlarian stage and the last zveno but two to the Sicilian - Tsvermagalian stage. The 200 ka climatic cycle caused mostly by the displacement of the ecliptic and clearly pronounced in high latitudes (Sharaf, 1974) can also be seen in Table 6.3. Fillon and Williams (1983) have analyzed the climatic contents of the 200 ka cycle. However, we know nothing of the attempts to relate these two cycles (400 ka and 200 ka cycles) t o the Pleistocene classification, although they are closely connected with this problem. It can be seen in Table 6.2 that the conventional transgressions and erosional unconformities of the Mediterranean are in better agreement with some parts of the 180-230 ka cycle than the traditional interglacials and glacials of northern Europe. The duration of the sybcycles of the 200 ka cycle in these two regions is however different. Therefore, the unification of the climatostratigraphic nomenclature on the basis of the Mediterranean, i.e. on the basis of the 200 ka cycle, would have been artificial for glacial regions. Association of the 400 ka cycle with the oxygen-isotope scale gives more advantage, since it allows us to solve the problem through the usage of global information and combination of palaeoclimatic and biostratigraphic evidence. Our proposal is to accept orthoclimathems and superclimathems as the major units of a unified climatochronoiogic classification of the Pleistocene. It can be specified here that an orthoclimathem (OCT) is understood as a range of sedimentary rocks that corresponds to the globally observed and regionally identified temperature trend imprinted in the changes (whatever form they take) of the local natural environment, whose length is no less than ten thousand years.2 The volume and successsion of orthoclimathems are determined by oxygen-isotope stages of the deep-sea Pleistocene. At the same time, since orthoclimathems are indentifed by the events that once took place, their boundaries and range should continually be ~ h e c k e d .Two ~ of the adjacent orthoclimathems correspond to a climatic cycle (rhythm) o f tens of thousands to a hundred thousand years long. In order for orthoclimathems to be broadly used in dissecting continental sediments, they should have parastratotypes in the coastal shelf facies. Parastratotypes should be selected in such a way that they can be reliably correlated with deep-sea sediments and isotope stages on the basis of micropalaeontological (as well as magnetostratigraphic and radiological) methods and, on the other hand, it is necessary that they can be divided into thermomers and kryomers, which are correlated with loess and glacial sequences with the help of the methods used in conventional climatostratigraphy and biostratigraphy, including magnetic and radiological techniques. The coastal sequences of the Black Sea and of the Mediterranean will suit best o f all as parastratotypes of orthoclimathems. Superclimathems (SCT) are suggested to designate a range of sedimentary rocks 'Since the amplitude of temperature variations at different latitudes can vary within different ranges of the geochronological scale, it is hardly expedient to include it into the definition. 30rthoclimathems can correspond to some groups of isotope stages (for instance, rtages 2 + 3 + 4 or 8 + 9 + 10) o r to substages (for instance, substage 5e). But in this case also it would be advitable to call orthoclimathems according to the nomenclature of the Shackleton - Opdyke isotope stages (1973, 1976).

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that corresponds to the globally observed and regionally identified complex kryomers and thermomers, which are related in pairs to the 400 ka climatic cycle. In the Pleistocene, superclimathems are represented by a group of three or more orthoclimathems, whose volume and composition is in the long run determined by agreement, but so that their boundary in each unit can be observed globally in the best possible way. The lower boundary of thermo-superclimathems (thermo-SCT) is considered to be the greatest (over a 400 ka interval) termination in the deep-sea isotope sequences and the beginning of the warmest (over the same interval) interglacial in the continental sequences. Such boundaries are transitional between isotope stages 6 and 5 as well as 16 and 15 (Terminations 111 and IX) and the relevant Eltigen - Mikulino and Patrai - Cromer interglacials. The lower boundary of kryo-superclimathems (kryo-SCT) is more conditional and vague. It can to some extent be observed with the changes in the organic world, for instance the appearance of periglacial complexes (Swanscombe - Volga - Khazar and Nistrus - Griice) and the main changes in the geomagnetic field. Thus, it is just this boundary that is characteristic of the Brunhes - Matuayama and Cobb Mountain reversals as well as the r6 Levantine excursion, the largest in the Brunhes orthomagnethems. The kryo- and thermo-superclimathems of the Pleistocene are of almost the same length from 180 to 220 ka. However, in the Pliocene their length can be different. It is interesting to note that superclimathems exhibit certain geological and geomorphological features. Thus, the continental thermo-SCT are characterized by prevailing alluvial and bog sediments, whereas geomorphological processes are characterized by deep erosion and the formation of deeply incised valleys. Along marine coasts thermo-SCT are distinguished by the formation of terraces and coral reefs at a level exceeding the present sea level. The continental kryo-SCTs chiefly exhibit other genetic types such as tills, loess, deluvium-solifluction areal sedimentation. The river valleys within the kryo-SCTs are filled with and buried under periglacial alluvium and deluvium as well as congelifluction-sors assemblage, till and loess deposits. In the coastal regions kryo-SCTs, including orthothermomers, are formed as a rule below the present sea level (although these terraces have sometimes been elevated even higher as a result of subsequent processes of emergences). The indicated geological and geomorphological differences themselves are indicative of considerable changes in the global climatic system within the 410 ka cycle. Therefore, it is not accidental that in West Siberia (Fig. 5.6) and the Don Basin the age of the present river valleys does not exceed the limits of the first superclimathem. And each subsequent thermo-superclimathem counting by its age is represented by the sediments of buried river valleys (the IIIrd SCT) and more ancient pre-river valleys (the Vth SCT).

(1) The number of Pleistocene climatic changes as fixed in the continental sequences of the shelf, loess and glacial assemblages is similar to the number of isotope stages in the deep-sea sediments.

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(2) In the Pleistocene, climatic changes proceeded synchronously in all areas under consideration; however, the response of the natural environment to these changes was not the same in all the places because of the different final rate of various geographical processes; (3) The terms “glacials” and “interglacials” turned out to be ambiguous and useless in correlating the climatic events in high and low latitudes. For example, in high latitudes the “interglacial units” are three to four times shorter than the glacial ones, whereas in the Mediterranean it is just the opposite. The oxygen-isotope variations as an integral reflection of fluctuations in the surface ice volume play to a certain extent the role of a general standard (“a ruler”) of the global climatic events in the Pleistocene, the orthoclimathems. (4) Orthodimathems are the globally observed synchronous temperature trends that last tens of thousands of years and can be identified regionally by a set of bio-, magneto- and radiological data. Their indices are the same as for the oxygen-isotope stages. (5) An analysis of the temporal principles governing the Pleistocene changes allows us to distinguish natural groups of orthoclimathems corresponding to the kryomer and thermomer portions of the 400 ka climatic cycle called superclimathems. These are the units of the global distribution (even for the Pliocene) identifiable by sufficiently pronounced local palaeontological characteristics.


Chapter 7

CLIMATIC CHANGES IN THE EARLY AND MIDDLE PLEISTOCENE

7.1. Introduction The factual information presented in Section I1 on climatic changes in the Pleistocene allows us to describe natural zones and climates of the Northern Hemisphere within all orthoclimathems. The framework of this book, however, does not allow us to describe all of the more than twenty recognized orthoclimathems. Moreover, the main purpose of this work is to elaborate the time structure of the past climates. The reconstruction of natural zones and climatic environments of the Pleistocene orthoclimathems will perhaps be done in future. Therefore, we present below a concise review of the natural environments and climates of the 12 regions in the Northern Hemisphere that have been considered earlier with an emphasis on the most important events. Since this review is going to be quite brief, we shall not overburden it with references, which have been given in Section 11.

7.2. The sixth (Giinz) kryo-superclimathem, 1.17 - 1.0 Ma The beginning of the sixth SCT conditionally called the Giinz superclimathem coincides with the base of the Sicilian stage in its type area in the vicinity of Palermo, where it is dated at 1.2 - 1.15 Ma (Ruggieri and Sprovieri, 1977; Ruggieri et al., 1984). Its prototype can be the double Menapian kryomer in the Netherlands (Van Hammen et al., 1971; Zagwijn et al., 1971; Zagwijn, 1975) and the Port Katon - Kvemonataneby) kryomer of the Ponto - Caspian Basin in the USSR. According to palaeomagnetic data, these kryomers are found below the Jaramillo boundary. In the sequences on Tsvermagal Mount this cold stage occurs within the normal polarity event (= Cobb Mountain ?, 1.1 Ma). In North America the standard of the sixth superclimathem is the Nebraska “B” moraine underlain with the Coleridge ash of the Pearlett S-type dated by the fission-track method at 1.27 Ma (Easterbrook and Boelstorff, 1981). In the Patagonian Andes it is correlated with the moraine of the greatest piedmont glacier with an age of about 1.2 Ma (Mercer, 1978) and in New Zealand it is associated with the kryomer containing the Pikihikura ash with fission-track date of 1.06 Ma (Hornibrook, 1981). In the deep-sea sequences the sixth SCT is fixed by the first appearance of small Gephyrocapsa of about 1.13 Ma in the roof of zone NN 15 (Rio, 1982) and Mesocena elfiptica of about 1 Ma (Berggren et al., 1980). In the DSDP key site 504 near the coasts of Ecuador it corresponds to the zone of the Mesocene cool peak, which falls between 1.3 and 1.0 Ma (Bukry, 1983). The correlation of the sixth superclimathem with the isotope curve cannot be accurate, because the number of isotope stages below stage 25 does not coincide in different authors.

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Climatic conditions of the sixth SCT can first of all be estimated by the data on the extent of glaciation. The advancement of the Laurentian ice sheet, which led to the formation of the B-type tills, reached as far as Nebraska. It cannot be excluded that the lower tills on Island Banks were also associated with that glaciation. In Europe, there is also indirect evidence of glacier advancement at that time at least as far as the Volga/Oka watershed. For instance, while examining the Middle Goryanka alluvial deposits in the Don Basin associated with the normal polarity zone (probably, the Cobb Mountain event), Krasnenkov discovered pebbles of the Scandinavian crystalline rocks washed out from the north through the river valley, which presumably (Fursikova, 1982) linked the upper reaches of the Volga and the Don. Thus, it can be seen that the Menap - Nebraska "B" - Gunz glaciation was not smaller but even greater than the Wurm glaciation. This conclusion has been corroborated by strong evidence that the Cassian - Calabrian regression dated at about 1.1 to 1.0 Ma was very extensive (Ambrosetti et al., 1972). At the time of this regression and the Iron Creek glaciation the Bering Strait was drained, which can be inferred from the appearance of the Asian migrants in North America (Archidiskodon haroldcooki) and the American migrants among the Oler fauna. The vestiges of periglacial climate of the sixth SCT have been preserved in the 11yichevsk loess in the Ukraine and Wucheng loess on the Ordos Plateau in China as well as in the loess sequences with Kuruksai fauna in Soviet central Asia and several stages of polygonal-vein ice in the lower Oler suite of north-eastern Asia containing fauna of the early Irvington type. The tundra steppe phyto- and zoocenose moves from Central Asia to central Europe, where it is represented by the Graze climatic zone (Chaline, 1977). Thus, it can be stated that the sixth superclimathem is an intense and evidently multi-phase global cooling accompanied by extensive advancement of ice sheets (as far as 41"N in North America and 56"N (?) in Europe) and loess formation. Clear traces of the lacustrine transgressions in the arid zone have also been preserved since this cold stage. Among them there are the Middle Apsheron transgression of the Caspian Sea and that of the lacustrine beds, unit F, in the section of site KM3 on Lake Sierls (Smith et al., 1983). Both transgressions are dated by the presence of normal polarity event n3 ( = Cobb Mountain ?).

7.3. The fifth (Giinz - Mindel) thermo-superclimathem, 1.0 - 0.76 Ma The standard of the fifth thermo-superclimathem is the Portuensian transgression and its Black Sea equivalent, the Tsvermagal transgression with stratotypes on Mount Tsvermagal and along the Chakhvata River. In the deep-sea sequences the fifth thermo-SCT corresponds to isotope stages 25 - 21. Along the Mediterranean coast the Sicilian deposits are incorporated into the third and fourth marine terraces lying at a height of 105 to 30 m above sea level (Zeuner, 1959, 1965; Kaiser, 1965); for example, that is level U on Mallorca according to Butzer (Bowen, 1978). In eastern central Baja California, they presumably correspond to high marine terraces near Santa Rosalia and San Climente, the earliest being estimated by Ortlieb at one million years old (The XZth ZNQUA Congress, Abstracts, vol. 11, p. 229).

21 1

In the Bering Sea the fifth thermo-SCT is marked by the Anvil warm-water transgression, at the time of which the Pacific molluscs and foraminifera easily invaded the Polar Basin and moved along the Canadian shelf as far as the Norwegian Basin (Herman and Hopkins, 1980; Gudina et al., 1984). At that time for many years pack ice did not exist in the Polar Basin, which means that the then Arctic climate was much warmer than nowadays. This is recorded quite well in the cores from the central Polar Basin (Clark, 1982; Herman, 1975). In the loess sequences, the fifth thermo-SCT is represented by three soil horizons: the soil L - K - J (Kukla, 1977; Fink and Kukla, 1977)-Nogai (Lebedeva, 1972) and the Shirokino pedocomplex (Veklich, 1982). In the Danube Basin and in the Ukraine they are all of subtropical type formed in a hot climate under seasonally humid conditions. The soils are associated with Banatica mollusc assemblages. In the European forest zone the fifth thermo-SCT is represented by three recently discovered interglacials. In the Netherlands it is the Bavel, Leerdam and Waardenburg thermochrons (Zagwijn, 1985), in West Germany it is the Waterson, Pinneberg and Osterholtz thermochrons (Menke and Behre, 1973). Their analogues are recognized in eastern Europe in the West Siberian Lowland. The spore and pollen record of all three therrnomers reveal forest, mainly broad-leaved assemblages containing a number of endemic forms such as Eucommia, growing now in southeastern Asia, and Tsuga requiring uniform precipitation. At that time the flora contained a large number of exotic species (over 25%). The fifth thermo-SCT is divided by two cold phases corresponding to isotope stages 24 and 22. They are represented by the Chumbur loess sequences in the Sea of Azov region, by the Linge and Dorst - Elba kryomers in north-western Europe and the Mediterranean Vallonnet and Ficarazzi kryomers. According to the data of Byelorussian (Voznyachuk, 1978, 1985) and Polish (Lindner, 1981, 1984) scientists, northern Europe witnessed at that time the formation of an ice sheet whose size was about the size of the Wurm glaciation. It seems likely that the lowest tills from the sequences in the vicinity of Moscow (Lika till), in the upper reaches of the Kama and in the Irtysh issue (Mansi till) were formed during this glaciation, although it is more probable that they developed later (during stage 20). The most ancient interglacial of the Russian Plain (Mikelevshchina) with Arucites johnstruppi and Brusenia bielorussica is more likely to correspond to isotope stage 21. The fifth SCT is associated with a certain mammalian complex, which is everywhere represented by a transient Tamanian - Tiraspolian fauna with the latest representative of the southern elephant (Archidiskodon meridionah enikendis) and the Pleistocene rodentia with Mimmomys rarricepoides (= M . oeconomus). Many authors recognize this fauna as an individual assemblage (Petropavlovka - KaraiDubina - Sent-Prest complex and so on). On the whole we may state that climatic environment throughout the fifth SCT was similar to the Late Pleistocene, i.e. warm interglacials with summer temperature exceeding the present one by 2 - 5°C alternated with short phases of cold and arid climate. Such frequent changes of climatic and consequently natural environment was observed in particular by the Byelorussian and Lithuanian palynologists in the sections of the Brest and Daumantai series (Makhnach et al., 1981; Kondratene, 1979). It is natural that frequent climatic fluctuations inducing shifts in the natural

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zones were favourable for fast evolution of mammalian fauna and our distant ancestor Homo erectus, who at that time migrated from Africa to Europe and Asia. In the Caspian Basin the fifth SCT corresponds to a long Tyurkyan - Duzdag regression, whose age determined by ash horizon “B” in Azerbaijan by the fissiontrack method is found to be 950 - 1,050 ka (Ganzey, 1984) and by the K/Ar method as 850 k 250 ka (Zubakov, 1974). The Tyurkyany sequences are evidently observed in the Jaramillo event. In the sequences of Lake Sierls, the Great Basin, the fifth superclimathem is evidently associated with unit “E” represented by alternating salt and lacustrine ooze deposits, which have also been found in the Jaramillo subzone (Smith et al., 1983).

7.4. The fourth (Mindel) kryo-superclimathem, 760 - 585 ka The fourth kryo-superclimathem is associated with three successive ice advances synchronous with isotope stages 20, 18 and 16 that left behind moraines in the central part of the Russian Plain, the Lika, Setun - Lipetsk and Pereksha- Don, three tills in Byelorussia, in Poland (Narew, Nida and San tills) and in North America (tills of types A). They correspond to three kryochrons in the Netherlandish sequences (“Glacials A - B - C”) and the Early Elsterian glaciation. The last ice advance was the greatest Pleistocene one in many regions of the Northern Hemisphere. That is why it was recognized in the Don Basin as the Dnieper one. Actually the Don

......1 --2 -

3 --4

-5

-~--6-7

-8

-1-9

Fig. 7.1. Inferred boundaries of glacial advances in western Eurasia. 1 - Narew-Unstrut (OCT 22 or 20?) and Nida-Narev-Kama-Mansi (OCT 20?). 2 San - Setun - Lipetsk (OCT 18), 3 - Don - Dzukiya - Mogielanka - Podkamennaya Tungusska (OCT 16), 4 - Elster- Wilga-Dainava-Oka (OCT 14 or 12?), 5 - Saale-Odra-Dnieper (Early Moscow) - Samarovo - Bakhta (OCT 10- S), 6 - Warthe - Late Moscow - Yenisei (OCT 6 ) , 7 Weichsel - Vistula - Valdai - Ermakovo (OCT 5d - 4 - 2), 9 - Dryas 111.

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and Dnieper moraines of the Russian Plain appeared to be rhythm analogues, i.e. they formed during the same climatic phase of a 400 ka rhythm but at different cycles (Fig. 7.1). The “Mindel” SCT is divided by two interglacials, when the climate was relatively cold. During the first (Ilyinka - Karchevo - Akulovo - Kozi-Grzbiet - Westerhoven) interglacial, central Europe was covered by mixed forests and inhabited by subtropical animals. For instance, porcupines (Hystrix) penetrated as far as Czechoslovakia and Byelorussia. The second (Moiseevo - Pilczyca) warming was more moderate and the summer temperature probably did not reach the present one. A deep Flaminian sea regression is K/Ar dated at 680 - 706 ka by Lacium ash and reached its maximum during isotope stage 18. Isotope stage 19 corresponds to the end of a relatively cold Portuensio transgression, when the Mediterranean was inhabited by Hyalinea baltica. It probably corresponds to the Tiltim - Bolgokhtoch Pinakul marine glacial layers with poor foram and ostracod fauna studied by Gudina (1969); Gudina et al., (1984), Slobodin and others (Danilov et al., 1984). In the Ponto - Caspian Basin the “Mindel” glaciation has long been acknowledged t o correspond t o a double Baku transgression, whose water got through the Manych Strait into the Black Sea depression and probably through the Bosporus Strait into the Aegean depression. The Baku or Platovo layers with Didacna parvula, D. rudis are folded into the upper portion of the Chauda - Baku stage, which had not been subdivided earlier. According to the fission-track method, the age of the Baku stage is between 500 and 700 ka. The analogue of the Baku transgression in the sequences of Lake Sierls appears t o be ooze of unit “D” dated by the Brunhes - Matuyama reversal (Smith et al., 1983). The “Mindel” kryo-SCT has everywhere a pronounced faunal characteristic: it is associated with a typical Tiraspolian - Galerian fauna including in addition to Marnrnuthus trogontherii many tundra-forest animals such as Ovibos, Rangifer, Dicrostorzyx, Lemmus and others (Alexandrova, 1976; Agadzhanyan and Erbaeva, 1985; Vangengeim, 1970). The entire complex of geological and palaeontological evidence shows that the climatic cooling of the fourth SCT was very strong. At the same time it is impossible t o fail to draw the conclusion that this period was more favourable for the expansion of ice sheets than orthoclimathems 24 - 22. That means that the climate of the fourth SCT was rather “snowy”. Does this only refer to the Northern Hemisphere? There is no direct evidence to compare the situation in the two hemispheres. However, taking account of the almost identical peaks of isotope stages 22 and 16 (as well as 12 and 6) and an undoubtedly great volume of ice in the Northern Hemisphere within stages 18- 16 (and 8-6) compared with stages 24 - 22 (and 14 - 12), we can reach the conclusion that the culminations of the surface glaciation in both hemispheres were asynchronous.

7.5. The third (“Mindel - Riss”) therrno-superclimathern, 585 - 350 ka Ancient alluvial strata of buried valleys known as Mariinsk and Strelitsa suites on the Don River, the Gunki suite on the Dnieper River and the Tobolian and Larjyak

214

suites in West Siberia, are the standard of the third thermo-SCT in the continental sections in the USSR territory. They include at least two interglacials: Muchkap and Likhvin on the Don River, Vorogovo and Panteleyev on the Yenisey, and Scorodum and Chembakchino on the Ob River. The lower part of these strata is characterized by the remains of the Late Tiraspolian mammalian fauna and at the same time by thermophylic Kolkotovian complex of freshwater molluscs with Viviparus tiraspofitanus (on the Dniester), Corbicula fluniinalis etc. The above interglacial of 370 - 500 ka, according to the thermoluminescene dating technique, which correlates with the 11 - 13th isotopic stages, contains mammalian fauna of the Singulian type with Palaeloxodon antiquus and Arvicola mosbachensis and Volgan type with Mammuthus chosaricus and A . chosaricus. It corresponds to the Likhvin interglacial of the Russian Plain and the Holstein of western Europe. At the same time the latter two names are used to designate the whole third SCT. Interglacial alluvial members are divided in the West Siberian sections by silting regional stages with tree stubs and forest-tundra pollen spectrum (Sarchikha kryomer = the 12th or 14th isotopic stages), which corresponds to the last findings of the Tiraspolian fauna. In Kukla’s scheme (1977) the equivalent of the Sarchikha kryomer is called Elster-11, and that of Vorogovo and Muchkap is called “Holstein - Frimmersdorf”. In the sections of drilling wells in the Russian Plain the third SCT is called the Odintsovo - Roslavl (Moskvitin, 1957) - Shklov (Goretsky, 1980) interglacial. It has three individual optima according to its spore-pollen characteristics: the lower Glasov ( = Lyubny, OCT 15c) with oak and elm peak, and the second Pepelovo ( = Lysogorsk, OCT 15a) with hornbeam peak, and the third Galich (OCT 13) with Likhvinian-like pollen diagram (Moskvitin, 1958, 1967, 1976; Goretsky, 1983; Yolovicheva, 1979; Chebotareva, 1984; The Moscowian ice sheer . . ., 1982). Since this triple interglacial lay on the moraine of maximum glaciation, i.e. was traditionally considered, on the Dnieperian, then the Odintsovo - Shklov triple interglacial was identified as the “second Middle Pleistocene”. Only Vosnyachuk (1965, 1978) assumed the Shklov - Roslavlian layers to be of Cromerian age. This idea was confirmed by Biryukov (Marginal Formations. . ., 1985, p. 106), who found the Tiraspolian rodent fauna with Mimomys intertnedius in the Glazov layers of the stratotype section near the city of Roslavl. Now we can name the entire third SCT the Odintsovian comparing it with isotopic stages 15c, 15a and 13, and two intermediate kryomers (Podrudnya and Oka) with stages 15b and 14. In Poland, the Great and Mazovian interglacials (Rozicky, 1969) were analogues of the Odintsovo thermo-superclimathem divided now into two interglacials - Ferdynandow and Barcowike-Mokre (Lindner, 1984; Moijski, 1985). In marine sections of the Mediterranean and the Black Sea, the third SCT includes the Tarquinian - Uzunlarian triple transgression (Table 6.1 ) separated by bipartial Nomentanan - Palaeoeuxin erosion phase, the equivalent of the Oka - Elster I1 glacial advance. In the section of marine sediments in the north of Eurasia, the Tarquinian is equivalent to the Kolva - Ust’ Solenaya layers with rather warm water for the Arctic complex of foraminifera, and the Riano - Padimei - Kochos layers with a cooler complex. All this agrees with the interpretation of the Mindel-Riss given by Penck and Bruckner (1909) and particularly by Beck (1934) and Eberl (1930), who revealed in

215

the “great” and long-term Mindel - Riss two interglacials: Kander and Gluch (Moskvitin, 1970), which now can be synchronized with isotopic stages 14 and 12. This supposition of the author (Zubakov, 1968c) has been confirmed now by considerably larger extensive data. In loess sequences, the Mindel - Riss (without inverted commas now) thermoSCT is represented by a triple (sometimes quadruple) pedocomplex called the Tsokur pedocomplex in the Asov Basin and the Gorodskoi pedocomplex in the Don Basin. In the Ukraine, it corresponds to the Martonosha, Lubny and Zavadovka soils with Acheulian flint tools. The latter serve as indicators of the soil of the third SCT in Asia and in the Mediterranean Basin (Terra Amata and others). Thus, the third SCT is a well-documented triple interglacial, corresponding to isotopic stages 15, 13 and 11. They differ distinctly by their palaeontological characteristics (the Late Tiraspolian, Singilian and Volgan mammalian complexes, Kolva and Padimei foraminifera complexes). During all these interglacials the climate was very warm with temperatures exceeding the modern one by 2 - 3 ° C (Kondratiene, 1977; Zagwijn, 1973, 1975; Makhnach et al., 1981), while the moisture content was different. The middle, Likhvin - Holstein, interglacial was characterized, judging by the predominance of dark coniferous forest, by tsuga, hornbeam and the presence of yew in the forests of the middle zone of Europe, by more humid, even marine, climate extending almost as far as the Urals. Winter was particularly warm with temperatures of 2 - 4°C higher than at present, which allowed subtropical plants such as vine (Vilis),yew and animals such as Hystrix, Macaca and Hippopotamus as well as Pdaeoloxodon antiquus to penetrate northwards of the present border. This enables one to conclude that the narrow time sections along the optimum of the Likhvin - Holstein (OCT 13) and Glasovo - Cromer s.str. (OCT 15) interglacials can be used for spatial palaeoclimatic reconstructions. In particular, the former, since it is quite recognizable in the deep-sea sections: it lies just under LAD Pseudomiliania lacunosa, 440 ka ago.

7.6. The second (Riss) kryo-superclimathem, 350 - 130 (170?) ka

The second kryo-SCT includes the sediment envelope with buried river valleys formed during the third superclimathem. It is represented by genetically unhomogeneous series: in the north by moraine and fluvio-glacial deposits that correspond to two advances of the ice sheets (Dnieper - Moscow, Saale I - I1 - Illinois I - 11) of which the second was maximum; in the periglacial zone, by lake sediments and periglacial formation. In the extraglacial zone, the equivalents of moraines in the interstream areas are two loess horizons (Orel and Dnieper in the Ukraine), and terrace levels on the Caspian Sea (two upper Khazarian terraces). These sediments are characterized by the presence of remains of the Volgan ( = Khazarian = Aldenian) mammalian fauna with Mammuthus chosaricus, M.primigenius fraasi, M .pr.pa vlovae, Coelodonta antiquifat is, Equus cabaIlus, Discrostonyx simplicior etc., which were distributed over great areas in Eurasia. The geomorphological criterion was important in dissecting the sediments of the

216

second superclimathem. In this case, of particular importance was the final advance o f the Warthe ice sheet, whose end-moraine belt occurred in the North Germany lowland and further eastwards to Moscow. Here, it was associated by some researchers with the moraine of the Kalinin glaciation (Moskvitin, 1967, 1976) and by the others with the Moscow glaciation (Markov et al., 1965). In West Siberia also, some researchers (Saks, 1948) correlated with the Warthe the Zyryan moraine, while others (Zemtsov and Shatsky, 1953), the Tazovian moraine. Unreliability of morphostratigraphic substantiation was fatal for the chronostratigraphy of the glacial formations of the Central Russian Plain and West Siberia. It appears now that this led t o an incorrect interpretation of borehole logs, in which for the most part the interglacial strata were revealed (frequently bedded in the erratics), and naturally to incorrect correlation of interglacial sequences. It turned out that Moskvitin (1967, 1976) was right and the Warthe moraine corresponds in reality to the Kalinin one in the Upper Volga Basin. However, its age proved to be older; according to thermoluminescence dates it is 140 - 200 ka and corresponds to isotopic stage 6 like the age of the Yenisey ( = “Zyryanka”?, Taz?) moraine in Siberia. Consequently, the Moscow moraine in the Upper Volga appeared to be an equivalent of the Dnieper one in the Chekalian type section, and the “Dnieper” moraine from the well sections in the Moscow region, of the Lower Pleistocene (the Don) one. Thus, there appeared great stratigraphic discrepancies. These were partly removed in 1984- 1985 due to investigations carried out by geologists from the Central Geological Survey (Krasnenkov et al., 1984; Shik, 1981; Marginal Formations . . ., 1985). Now it is clear that the time extent of the Moscow glaciation is greater than that of the entire second kryo-SCT. The Moscow glacial complex of the upper reaches of the Volga includes the Kaluga OCT corresponding to the Liwiec stage in Poland and Fuhne in GDR (isotope stage 12 or lo), the Dnieper OCT land and Fuhne in GDR (isotope stage 12 or lo), the Dnieper OCT ( = the Warthe, stage 6). It is possible that there is one more intermediate glacial advance corresponding to isotope substage 7b. These orthoclimathems are separated by the Chekalin - Domnitz - Zbojno - Hoogevin (stage 11 or 9) thermomer and the Cheremoshnik “B” - Grabowka - Treene ( = stage 7) thermomer. The first has thermoluminescence dates of 300 - 340 ka and the second 180 - 245 ka. A similar situation occurs in North America, where the Illinois complex is traditionally considered (Reed et al., 1965) to consist of three horizons : Liman, Monican and Buffalo-Hart separated by three horizons of weathering (humbotill). Definite confirmation of two interglacials within the second SCT is a series of dates of speleothems from the caves of the Rocky Mountains in Canada, revealing two nonglacial intervals between 215 and 320 ka and between 185 and 235 k a (Harmon et al., 1979). These estimates correspond t o the thermoluminescence dates for intraRiss thermomers in the USSR and in Poland. In the deep-sea section the second SCT is synchronized with isotopic stages 10 - 8 (or 12-8?). The maximum shift of 6180 is associated with stage 6, whereas the greatest extent of the Northern Hemisphere glaciation and the greatest (Ostian) regression correspond to stage 8. From this we can conclude that either the Antarctic ice sheet was much more extensive within stage 6 or, as Williams et al. (1981)

217

believe, within the same stage there formed in the Arctic a sea ice cover more than 1 km thick resting on the sea bottom. There is no reliable information about the climate within stages 11, 9 and 7 . After Zagwijn (1973), the temperature of the warmest summer months during the Hoogevin thermomer (stage 1 l ? ) was 1 "C lower than the present one, and during the Bantega interstadial (stage 9?) it was 3°C lower (Fig. 5.2).

RCsumC (1) For the last three million years of geological history the most distinct climatic and at the same time biostratigraphic boundary is the end of the Villafranchian in the continental sequences and the onset of the Sicilian transgression in the Mediterranean. This boundary coincides with the first continental glaciation of Eurasia and the erosional phase, with which the formation of the modern river systems begins. This boundary is everywhere dated at 1 f 0.1 Ma. (2) The time structure of the Early and Middle Pleistocene climate (1.17 - 245 ka BP) is much more complicated than was thought earlier. During that time there were no less than 10 t o 12 global coolings accompanied with the advances of ice sheets far into the middle latitudes and the same number of global warmings accompanied with the glacioeustatic transgressions. (3) The global observation of all orthoclimathems is frought with difficulties. Therefore, only superclimathems can be recognized as inter-regional units of the Early and Middle Pleistocene. There are five of them and it is advisable to give them, through international agreement, the Alpine nomenclature.


Chapter 8 CLIMATIC CHANGES IN THE LATE PLEISTOCENE

8.1. Tyrrhenian ( = Riss - Wurm sensu lato) megathermochron, 245 - 118 ka 8.1.1. On the time-scope of the ‘‘Riss- Wiirtn” (277-244 ka)

Table 6.2 shows that the time-scope of the Riss - Wiirm interglacial adopted in the Mediterranean and glacial Pleistocene systems differs considerably. In the Mediterranean system it embraces the entire horizon with Strombus fauna, which corresponds, by radiological dates (220 ka to 85 ka), to isotope stages 7, 6 and 5. In the glacial system it includes a brief interglacial period in northern Europe between the Warthe and Vistula glacial stages equivalent to isotope substage 5e (127- 117 ka). Its shelf stratotype is considered to be the Fjmanger layers in Norway with U-series, TL and amino-acid dates between 130 and 115 ka (Mangerud, 1981; Miller et al., 1983) and the continental stratotypes are the Taubach travertines (see Section 5.2.). The European Riss - Wurm has always been regarded as synonymous with the Eemian interglacial of western Europe and the Mikulino interglacial of eastern Europe. At present we cannot, however, use these terms as a unified nomenclature since the extent of the Eemian and Mikuline thermomers in time is fairly vague. Unfortunately, the stratotypes of these widely-used Pleistocene subdivisions (the sequences of the Eem river near Amersfoort in the Netherlands and near the village of Mikulino in the Smolensk region, USSR) have been studied neither radiologically nor palaeomagnetically. The correlation of the climatic optimum of the Mikulino interglacial with the maximum of Eemian transgression is based on pollen evidence. Thus, Grichuk (Gerasimov and Velichko, 1982) considers the hornbeam peak in pollen diagrams of zone “m6” in Mikulino and zone “f” in the Eem as a reference marker for synchronizing these subdivisions. Although this correlation can be argued, for example, Selle thinks that the peaks of the oak and the nut-tree rather than hornbeam correspond to the Eemian optimum (see table 2.4 in Bowen, 1978), it nevertheless would have been acceptable provided only one interglacial took place between the Dnieper and Wiirm glacial, of which we cannot be sure. Thus, Sutcliffe (1986), Bowen (1978) and Shotton et al. (1983) have shown that the English equivalent of the Eem, namely the Ipswich, includes layers containing various mammalian fauna, which is possibly of different age. The pollen diagram of the Ipswich also exhibits diverse elements (West, 1968). The radiological age of the Ipswich is between 114- 135 ka and 174 ka (Gaskoyn et al., 1981). Therefore, Bowen (1978) finally distinguishes three Ipswich thermochrons: ( I ) the IpswichTrafalgar Square thermochron comparable with substage 5e, (2) the IpswichBrandon thermochron (174 ka) corresponding to stage 7, and (3) the Ipswich-Ilford thermochron equivalent to stage 9. Still earlier Kukla (1977) came to the same conclusion concerning the continental areas of north-western Europe, and recognized

220

the Eem-Skaerumhede ( = substage 5a), Eem-Ehringsdorf ( = stage 7) and EemEem (stage 9). The known Ehringsdorf sequences which have always been regarded as the continental parastratotype of the Eem (Zeuner, 1959; Woldstedt, 1958), the upper and lower travertine horizons containing interglacial flora and fauna with interbedded thin loess deposits (Pariser) are also found by U-series to belong to different thermomers. The age of the upper horizon is 115 - 118 ka and the lower one 205 - 162 ka, these radiologic estimates being in accordance with palaeontological dates. The Arvicola teeth in the lower travertines are more ancient than in the upper ones (Heinrich, 1982). As long ago as 1970 the author discussed the problem of different ages of the beds with Cyprina islandica of the Boreal transgression in the north of Western Siberia. It turned out that the dates by U-series method for the shells of Cyprina islandica from the village of Pupkovo on the Yenisey river (233 -t 10 ka) and from the Agapa river, 164 f 5 ka (Zubakov, 1972b) correspond to isotope stage 7 rather than 5 . The supposition of Ostrovsky (1974) about different ages of the Karangat Black Sea layers has now been substantiated by TL and palaeomagnetic data. And, finally, it is hardly possible to reject the fact that three Mediterranean sea levels with Strornbus fauna are of different ages (see Table 3.1). All of them belong to the Riss - Wiirm period. Out of the sequences in the Russian plain corresponding to the Mikulino interglacial, only a few have been dated by the thermoluminescence method. Unfortunately the assignment of the Mikulino age can be argued in each of them. Thus, the peat layers from the Cheremoshnik section with TL dates of 220-240 ka that are Mikulino according to Sukachev et al. (1965) are recognized as Odintsovo by Grichuk (1981) and Sudakova (1974). In the subsequent publications (Sudakova et al., 1981), the ravine sections in Cheremoshnik can be divided into two types: A and B. Gyttja of section A refer undoubtedly to the Mikulino interglacial, whereas these of section B, which include a number of exotic elements that are absent in section A (Picea s.Picea, P.obovata) are probably more ancient. The layers with alluvial peat in the sections from Lake Nero and the former Lake Tatishchevo TL dated at 95 - 110 ka that are Mikulino according to Sudakova et al., 1977) have been identified as Mologosheksna by Moskvitin (1967). The lower soil in the stratigraphic holes in Kostyenki with the Blake event in the roof and TL date of 170 ~ t _ 30 ka from the underlying load has been recognized as the Mikulino by Lazukov (1980) and the Early Wurmian by Spiridonova (Praslov and Rogachev, 1982). All these facts show that the problem of the time-scope of the Riss - Wurm, to be more exact of the number of Riss - Wiirm interglacials, has not yet been solved. That means that the Riss - Wurm palaeoclimatic reconstructions based on only palynological correlation (i.e. on the assumption of a single interglacial) which has been the practice before (Frenzel, 1967; Gerasimov and Velichko, 1982), turned out to be chronologically unreliable. According to our scheme, the first superclimathem (SCT) includes isotope stages 7 to 1. Within this subdivision there are three thermomers: the Early and Late Riss - Wurm and the Holocene, which are delimited by two kryomers. The lower boundary of the first thermo-SCT is considered to be the appearance of the tropical

22 1

Senegal fauna of molluscs such as Stron7bus bubonius, Conus quanaicus, Mytilus senegalensis and others in the Mediterranean Sea. In the deep-sea sequences this boundary is marked by the first appearance datum (FAD) of Emiliania huxleyi, which is dated from 275 ka in the ocean to 225 ka in the Mediterranean (Rio, 1982) and the last appearance datum (LAD) of Globoquadrinapseudofoliata dated at 220 ka (Berggren et al., 1980). In the continental sequences of Ehringsdorf and Skurlat the position of this boundary is probably marked by the appearance of Palaeoloxodon antiquus germanicus Stef.

8.1.2. The Early Riss - Wiirrn ka

-

the seventh thermo-orthoclimathem, 245 - 190

The stratotype of the seventh thermochron appears to be the first level with the Strombus fauna of the Italian sequences, cycle X , - , , of the Mallorca Peninsula and the Tobechik and Zavetnino beds of the Karangatian in the Black Sea. They 10 ka (Ambrosetti et al., 1972; Bowen, all are dated between 220 ka and 190 1978; this study). The indicated step correctly chosen by Zeuner (1959) and Yakovlev (1956) as the boundary between the Middle and Late Pleistocene palaeoclimatically exhibits a major change in the dominant tendency of the evolution of the Pleistocene glaciation. During the kryochrons in the time preceding this boundary, the volume and area of ice sheets increased and after it they gradually decreased. In the Arctic shelf seas the seventh orthoclimathem comprised the first Boreal transgression with the U-series age of 233 - 170 ka, i.e. the Kazantsevo - Pupkovo transgression of the Yenisey river, the Mechigmen one of the Chukotski Peninsula, the Kotzebu of Alaska. By the system of Feyling-Hanssen (1976), it evidently corresponds to the Cibicides teretis zone, although it has later been re-dated as more ancient on the basis of amino-acid racemization (Feyling-Hanssen et al., 1982). It can be seen in Table 6.1 that the thermomer of about 180-230 ka has been recognized in almost all the twelve regions under consideration. In the loess sequences it is represented by a full profile of buried soils, which often contain flinty tools of the Late Acheulian type. In the sections of glacial and periglacial zones this thermomer is considered to be reliable (the sequences of Cheremoshnik, Treene, Grabowka, Ehringsdorf, Illford, Shirta and so on). However, its age has remained vague until now. Therefore, it is necessary to conduct additional complex investigations in order to find a generally accepted standard of the seventh orthoclimathem in the glacial and periglacial sequences. The only thing that can be stated without any doubt is that the thermomers of Shklov, Roslavl and Polnoye Lapino in the USSR do not refer to this orthoclimathem. The seventh orthoclimathem is a double interglacial. This is clearly indicated by the division of the Lower Kerangatian horizon into two members of the marine sediments, the Tobechik and Zavetnino, of cycle X (Strornbus I) at two terrace levels (X, and X, and of the lower Riss - Wurm pedocomplexes into two soils as well as by the presence of two thermomer members in the Schleswig-Holstein sequences (Riigen and Treene in Cepek and Erd, 1982; Stremme, 1982). Particularly valuable is information from the Bermudas, which are a tectonically stable platform and

222

therefore constitute an ideal “tide gauge” for studying the position of the sea level in the past. In this locality a group of scientists gathered unique information on both of the Riss - Wiirm interglacials (Harmon et al., 1983). Within the seventh orthoclimathem, two marine terraces were formed along the Bermudas coast (Fig. 8.1). The most ancient level, the pre-Belmont formation dated at about 230 - 270 ka, was revealed at a depth of 12- 15 metres below the sea level (substage 7c). A younger level, the Belmont formation, which is as high as two meters above the sea level, was dated by corals Siderasfrea, on the basis of U-series at 200-228 ka (substage 7a). This level is separated form the first one by a sea level decrease of more than 20 m at a time interval of about 220 ka (Harmon et al., 1983). This information coincides with that obtained on Barbados (Bender et al., 1979), in New Guinea (Aharon, 1984, 1985), over the Kanto Peninsula in Japan (Mashida, 1975), in New Zealand (Pilians, 1983) and in California (Muhs and Szabo, 1982), which can be seen in Table 8.1. For the last 250 ka only three times was the sea level higher than the present one, namely during the climatic optimum of the Holocene and within substages 5e and 7a. That means that at the time of these three optima of the first SCT, the volume of land-locked ice was smaller than at present and the climate was consequently warmer. This is supported by the evidence of intrusion of Boreal - Luzitan fauna of molluscs and foraminifera up to the Yenisey river at Pupkovo - Timan time (see Section 5.6). After Troitsky (1979) the Pupkovo transgression dated by the U-series method at 233 k 10 ka is just the Kazantsevo transgression. The statistical ecological analysis of the faunal composition of marine molluscs from the Kazantsevo layers gave Troitsky the evidence that at that time the water temperature in the lower reaches of the Yenisey was 1 - 2°C higher and the air temperature was 2 - 3°C higher than at present. According t o Vigers’s data for lower travertines from the Bilzingsleben and Ehringsdorf sections in Europe, the mean annual temperature was

4 8 0‘ %, PDB

VZB-238

Age, ka

Fig. 8.1. Late Pleistocene fluctuations for Bermuda based upon the U-series and amino-acid racemization ages as well as the geological reinterpretation. Also shown is the deep-sea foraminifera1 ‘*0/’60 curve of core V 2 8 - 2 3 8 of Shackleton and Opdyke (1973) and the lithologic units which correlate to the documented sea level events. Modified from Harmon et al. (1983), fig. 6).

Table 8.1. Marine terrace levels (height, rn, in bracket

-

reconstructed

~

of sea levels, age, ka)

:I I

1 E

1

i

223

North

SSTs

V23-42 V27-20 9

5 5

0 25

50

rA

I

>-

75

0

9

13

13 6

K708-7

V27-116 9

5 10

14

("C)

18

13

South

K708-1 6 -

V29-179

8

17

10 .

,

'

22

18

14 '

'

16

12 '

I

N P N

V30-97 20

9

13

17

21

0 25 t.C\

75

co

0 100

I

.-f

12s

125

150

150

I75

175

2oc

200

22:

225

25C

250

c

a,

m

Q

Fig. 8.2. Estimated August sea-surface temperature based on species counts in seven subpolar North Atlantic cores (see Fig. 1.3, 62" - 40"N). Estimates are constrained to values greater than 6 ° C because of transfer-function limitations. Plots are stretched to time axis. (From Ruddiman and McIntyre, 1984).

225

1.2 - 2°C higher than the present one and there were no frosts in wintertime (Glazek et al., 1976). The calculations of Loiek and Srnolikova (Kukla, 1977) for the fourth sub-complex of the loess sequences in Csechoslovakia (Fig. 4.3) have shown that the mean annual temperature was 12- 13°C. In the North Atlantic cores, the summer surface water temperature was not lower than the present one (Fig. 8.2). It follows from the oxygen-isotope curves that the cooling of the seventh orthoclimathem was quite strong and should have been accompanied by a glaciation of the Late Drias type. The glaciation could hardly be greater because otherwise it would have been reflected in the growth of speleothems of the Rocky Mountains, which were continually increasing from 234 to 185 ka (Harmon et al., 1977). 8.1.3. The sixth kryo-orthoclimathem, 190 - 127 ka A long dispute between Woldstedt (1954 - 1958) and Zeuner (1959) that lasted for many years about the Warthe glacial stage and Eemian transgression seems to have been resolved in favour of a third alternative, which suggests that the Warthe was within the Eem (Kukla, 1977; Bowen, 1978). The author thinks that the same situation also occured on the Upper Volga. Here, according to Moskvitin (1967), the analogue of the Warthe stade, i.e. the end-moraine belt of the Kalinin glaciation, appears to be between two interglacial sequences of the Mikulino type. Let us conditionally call the upper sequence the “Mikulino - Mologosheksna” and the lower one with gyttja “B” of Cheremoshnik and Dolgopolka, the “Mikulino - OdintS O V O ” . ~ Judging by thermoluminiscence dating, one of them is twice as old as the other (see Table 5.1). This “Warthe - Kalinin” ( = Moscow) moraine with TL age between 127 and 230 ka extends into West Siberia (Fig. 7.1), where it corresponds to the Yenisey - Belogorie moraine with TL dates from 1 10 to 240 ka. Some scientists consider it to correspond to the Zyryanka glaciation (Troitsky, 1979; Arkhipov et al., 1977; Astakhov et al., 1986), and others to the final stage of the Samarovo glaciation (Lazukov, 1970 - 1972). According to the Polish researchers, the Warthe glaciation sea level is represented by two stadia1 belts, namely the Warthe s.str. and Wcra (Lindner, 1984). The same picture has also been revealed by the author on the Yenisey, where the main stade of the Yenisey glaciation is separated from the second stade, the Lower Tunguska, by the Strelnaya interstadial, when the climate was cooler than at present (Zubakov, 1972a,b). In North America, the Warthe - Yenisey glaciation is correlated with kryomers of the Nom river (Hopkins et al., 1967) and the terminal stade of the Illinoian (Andrews et al., 1984). The sixth superclimathem has also analogues in a11 loess sequences. In the type section of the Kerch Strait (the Sea of Azov) the sixth superclimathem is correlated with the kryomer of Geroevskoye I1 represented by sand dunes alternating with the Tyasmin loess marked by Levalloisian-Early Mousterian artifacts (Lebedeva, 1972). At that time the Caspian Sea run-off existed in the Manych

‘

According t o Sukachev (1954), all Mikulino peats of Cheremoshnik are within the Kalinin tills in the form of erratics (Moskvitin, 1967, p. 86).

226

Valley, of which the Early Girkan fauna with Didacna cristata is indicative (Popov, 1983). The Bosporus discharge evidently took place from the Black Sea into the Mediterranean, which can be seen from the “cold” sapropel 56 in the eastern basin dated at about 176 ka (see Section 3.4). 8.1.4. The Late Riss - Wiirm, or thermo-orthoclimathem 5e, 127 ( I 70?) - 117 ka Since the true age of the Mikulino and Eemian thermomers in their stratotypes is uncertain, we are to determine anew the Late Riss - Wiirm orthoclimathem as an analogue of isotope substage 5e. While doing so, the most important thing is to take account of the Mediterranean sequences, where substage 5e corresponds to the middle Strombus level that is hypsometrically the highest (9- 15 m) or cycle Y, on Mallorca with an age of 125 k 10 ka to 127 k 13 ka. Its Black Sea analogue is deposits of terrace I1 with Cardium tuberculatum, i.e. the Eltigen thermomer with a numerical age of 120- 129 ka. In north-western Europe the standard of the fifth orthoclimathem can be the marine Fj~sa nge rbeds with Liftorina litorea and Parvicardium (Mangerud et al., 1981) and their equivalents, the upper travertines of Ehringsdorf and Taubach dated at 90- 133 ka as well as the Bobbitshole layers in England with the Hippopotamus fauna (Shotton, 1983; Sutcliffe, 1986). In the USSR the equivalent of orthoclimathem 5e appears to be the interglacial watershed peats in the Kalinin and Yaroslavl regions containing pollen diagrams of the Mikulino type that are not covered with tills. According to Moskvitin (1967), they are the Mologosheksna interglacial layers covering the Kalinin tills, while according to Kozlov (Marginal Formations . . ., 1985, p. 142), they are the Mikulino layers overlying the Late Dnieper ( = Kalinin - Warthe) tills. In the Arctic and sub-Arctic, orthoclimathem 5e evidently includes the marine layers with the most warm-water Pleistocene transgression (Karginski - Ponoi - Shchuchya), a zone with Islandiella islandica and Trifarina angulosa, which until now have been considered to belong to intra- Wurm time according to the I4C-series dates. The first U-series dates for the Ponoi moluscs from the Kola Peninsula sections (86 - 114 ka) have shown that I4C estimates have been erroneous (Arslanov et al., 1981). In the eastern Arctic, orthoclimathem 5e includes the Valkatlen and Pelukian transgressions with the same U-series date of 100 ka (Hopkins et al., 1973). In loess sequences, orthoclimathem 5e comprises the upper interglacial soil marked by flint tools of the Early Mousterian type (Priluki - Salyn soil, cycle B,, MB and others), which is pronounced more clearly than the two underlying soils. The Late Riss - Wiirm orthoclimathem has not everywhere the same time extent. For example, in the loess sequences the terminal interglacial soil has TL dates between 170 ka and 100 ka, whereas the peak of the glacial eustatic transgression is everywhere dated at 130- 120 ka. How can this be explained? Interesting information has been obtained by Polish geologists in the sections of the Eemian transgression in the mouth of the Vistula. Two Eemian sequences have been identified in this region: the lower one containing a relatively poor fauna called the Sztum layers and the upper one with rich fauna, the Tychnowy layers, corresponding to the climatic optimum of the interglacial (Makowska, 1982). Since the lower layers are thin and contain cold-water fauna, they can not be correlated with isotope

227

stage 7 (about 170 ka). It is possible that the Sztum layers correspond to substage 6b?. In the Bermudas sequences orthoclimathem 5e corresponds to the greatest elevation of the ocean for the last 250 ka, which led to the formation of the Devonshire terrace with a height of 5 k 1 m. Eleven dates based on the U-series technique have been obtained there for corals, ranging from 118 k 6 ka to 134 k 8 ka with an average of 125 f 4 ka (Harmon et al., 1983). This estimate agrees well with the age of this level in other areas, namely on Barbados, New Guinea (Fig. 8.3), in California and so on (Table 8.1). At the same time on the Bermudas, a submarine terrace about 150 ka old has been revealed at a depth of 5 to 10 metres below the sea level. Thus, it seems that although the climatic optimum of the last interglacial occurred between 127 and 120 ka, the process of warming (and of the formation of interglacial soils) started to develop earlier, about 170- 150 ka BP, i.e. within isotope substage 6b.

8.2. Spatial climate reconstructions for the temperature optimum of the last therrnochron (isotopic substage Se), 125 - 120 ka The last interglacial is of special interest for palaeoclimatology. A group of researchers from the Institute of Geography of the Academy of Sciences of the USSR have recently carried out reconstructions of summer (July) and winter (January) air temperatures and annual precipitation for western Europe and the USSR (Velichko et al., 1982, 1983) based on pollen analysis by using a method proREF Vila

Vllb

Vl

PHASE V

IN NEW IV

GUINEA 1110 lllb /I

I

0

Fig. 8.3. Comparison of 6I8O coral reefs and deep-sea benthics from Meteor core 12392 (Shackleton, 1977). The isotopic sequence are plotted on the same scale of change from present values and are adlusted to yield the best agreement for the modern samples. Deep-sea 6 ' * 0 data are adjusted 10 match the chronology of the coral reef terraces (after Aharon, 1985, fig. 9).

228

posed by Szafer (1 954), Iversen (1973) and developed by Grichuk (1 978). By the data on the surface water temperature in the North Atlantic obtained through factor analysis of planktonic foraminifera (Barash, 1983; Nikolaev and Blyum, 1985; Blyum, 1982), air temperature charts for July and January have been built up for the larger portion of the Northern Hemisphere (Velichko et al., 1984). In Figs. 8.4 and 8.5 schemes are presented that have been compiled by the author on the basis of combining the above information with the CLIMAP (1984) and palaeobotanical data for the USA and Canada (Wright and Frey, 1965), for Great Britain (Shotton, 1978), for western and eastern Siberia (Volkova, 1977) and for western Europe (Frenzel, 1967). Interpolation of data was applied because of lack of palaeobotanical information on the following regions: the larger part of the USA, subtropical and tropical Africa, North America and Asia. Mean latitudinal temperature departures from modern values have been obtained by averaging data in the 5" latitude x 10" longitude geographical grid points (Table 8.2). The values of the Northern Hemisphere global temperature for January (February) and July (August) were calculated taking into account the areas of

Fig. 8.4. Deviation of the summer (July-August) air temperature from present time for OCT 5e. 125 - 120 ka BP.

229

latitudinal zones by the formula:

AT,,

=

EAT, cos p, c cos p; '

-

where AT,, is the mean temperature of the Northern Hemisphere, ATi the mean latitudinal temperature and pi the geographical latitude. During this optimum the air temperatures for the entire Northern Hemisphere were 1.6"C above the modern for summer, 2.4"C for winter and 2.0"C on the average for the year. Analysis of data in Figs. 8.4, 8.5 and Table 8.2 showed that the largest values of warming were recorded in winter in high latitudes, north of 50"N. In particular, on the Taimyr Peninsula, in the north-east of the USSR and in the north of Canada the air temperature was 10- 12°C above the modern. In the European territory of the USSR temperatures increased by 3-5"C, whereas in western Europe the

30

60

Fig. 8.5. Deviation of the winter (January - February) air temperature from present time for OCT Se, I25 - 120 ka BP.

230

Table 8.2. Mean latitudinal temperature differences between the optimum of the Late Riss- Wiirm orthoclimathem (5e) and modern epoch Latitude (day) T("C)

90 - 80 80 - 70 70 - 60 60- 50 50 - 40 40 - 30 30 - 20 20- 10 10 - 0

July - August January-February

7.6 8.0

6.0 7.4

4.8 6.5

3.8 4.7

1.6 2.4

0.3 1.2

-0.2 0.2

-

-

-

-

Mean global values I .6 2.4

temperature changed a Iittle, by not more than I - 2°C. A small drop in temperature took place near Iceland, which was probably associated with peculiar circulation in the Norwegian and Greenland Sea. In the South Atlantic, near the western coast of Africa and in the region of the Panama Isthmus water surface temperatures were 1 - 2°C below the modern, whereas in the Pacific a stable warming was recorded. Summer temperatures in high latitudes (except for northern Scandinavia and the North Atlantic) increased by 6 - 7°C. Over the larger part of western and eastern Siberia temperatures grew by 3 - 4"C, in the north-east of the USSR, north of Jakutia and on the Taimyr the warming reached 5 - 6°C. In the European territory of the USSR and in western Europe temperatures were elevated by I-2°C. A certain decrease in temperature (by 1 - 2°C) was observed over the territory of central Asia, in northern Arabia and Africa, which was evidently due to improved moisture conditions in these regions. The climate of the last interglacial was less continental and the temperature gradient from the pole to the equator decreased. This changed the circulation and moisture conditions in various latitudinal zones. Fig. 8.6 depicts reconstructions of annual precipitation (in departures from the modern values) for the temperature optimum of the last interglacial. The reconstruction of precipitation field was based on the relationship between moinstening coefficient (precipitation/evaporation) and sums of air temperatures above 5°C for different types of natural zones. The relationship has been obtained by Savina and Khotinsky (1984) for present conditions and used as a basis for reconstructing precipitation for different time intervals of the Holocene. For calculating evaporation we have used its dependence on mean annual air temperatures for different types of natural zones derived from data of the Atlas of Heat Balance. A vegetation chart for the Mikulino interglacial from Gerasimov and Velichko (1982) has been used for reconstructing natural zones. As can be seen from Fig. 8.6, with global warming higher than 1"C, almost all regions of the Northern Hemisphere, except for a small area in the Mediterranean, received more precipitation than at present. In the north of western Europe and European part of the USSR, on Taimyr, Chukotka, and in the north of Canada annual precipitation sums increased by 200 - 300 mm, i.e. by 50% compared with the modern values. The moisture conditions improved considerably in the south of the European USSR, central Asia, the trans-Caspian, subtropical Africa (Sahel, Sahara) and India. In reconstructing precipitation for subtropical (monsoon) regions of Africa and Asia, which are most poorly covered with pollen and other proxy data, we have

23 1

30

63

Fig. 8 . 6 . Deviation of mean annual precipitation (mm) from present lave1 for OCT 5e

taken into account the fact that during all the Pleistocene thermochrons the intensity of monsoon circulation increased, which is associated with seasonal redistribution of solar radiation due t o astronomical factors (Rossignol-Strick, 1985). The warmest isotope stages (5e, 7, 9 and 1 1 ) corresponded to the most favourable moisture conditions in the Sahara, Arabia, monsoon areas of India, Asia and Australia. The reconstruction of precipitation by Kershaw (1978) by pollen d a t a for the north-east of Australia (20" S) showed that improving moisture conditions in this region were recorded during a warming within isotope stage 5 (5e, 5c and 5a), whereas during coolings precipitation reduced and percentage of sclerophyll taxa increased.

8.3. The Wurm megakryochron, 1 1 7 - 15 ka 8.3.1. On chronological models of the last glaciation An example of the last glaciation, the time interval that is closest to us, clearly shows that palaeoclimatic reconstructions depend wholly on the existing stratigraphic chronometric schemes. I t is well known that there are two competing

232

timescales of the Wurm megakryochron, one of which is short and the other long. It was supposed at the beginning of this century, when all chronological schemes were based on the ideas of Milankovich (1930), that there were three stages in the development of the last glaciation with an age of 115, 72 and 22 ka (Eberl, 1930; Yakovlev, 1956; Zeuner, 1959). Introduction into practice of the radiocarbon dating technique has induced extensive studies concerned with the last pages of climatic history. As a result, a vast bulk of chronometric information has been accumulated on climatic fluctuations over the last 70 ka. This allowed the most ancient Wurm interstadials, the Amersfoort and Brorup, t o be dated most accurately on the basis of 14C by thermal diffusive isotopic enrichment at 68,200 k 1,100 and 64,400 k 800 years, respectively (Grootes, 1978) and the Saint Pierre interstadial, at 74,700 years (Stuiver et al., 1978). These estimates together with hundreds of final series I4C dates obtained for both the Wiirm and more ancient deposits, have formed a basis for developing the so-called short scale of the Wiirm. This scale is based on information from Wright (1961), Dreimanis and Raukas (1973), Kind (1974), Grootes (1978), Stuiver et al. (1978), Serebryanny et al. (1981), Zarrina and Krasnov (1984). These authors correlate the Wurm glaciation with isotope stages 4 - 2 and the Riss-Wurm interglacial with isotope stage 5. Some of the authors divide this interglaciation into two thermochrons: The Eemian S . S . and Saint Germain (Woillard and Mook, 1981) or the Mikulino and Odintsovo (Zarrina and Krasnov, 1984). Another distinguishing feature of the short timescale of the Late Pleistocene is an accepted interglacial within the Wurm period, for a suitable standard of which have been taken the Karginski - Ponoi marine sequences in northern Eurosia (Kind, 1974; Gudina, 1978), the Karukula beds in eastern Europe (Serebryanny et al., 1981), and the Karangat beds in the Black Sea (Semenenko et al., 1973). All these authors dated these beds by final I4C series at between 26 and 48 ka BP. The rehabilitation of the long scale of the Wurm and Riss - Wurm have been started with introduction of U-series dates for the coral reefs on Barbados and the development of the oxygen-isotope scale of the deep-sea Pleistocene. The work was first done by Broecker et al. (1968), Morner (1972) and Zubakov (1974), (table 64). The long scale of the Wurm and Riss - Wurm shown in Table 8.3 is based on new information presented in Section 11. The most important data have been taken from Arslanov (1984), Andrews et al. (1984), Mangerud et al. (1981), Menke and Tynni (1984), Miller et al. (1983), Woillard and Mook (1981). Of particular value have been all TL dates of the Late Pleistocene obtained in the USSR (Geochronology of rhe Quaternary . . ., 1985), Poland (Marusczak et al., 1983; Lindner, 1984) and in France (Wintle et al., 1984). The suggested version of the Wurm and Riss-Wurm long scale has been developed proceeding from two criteria for checking the reliability of estimates, which are called below geological and chronometric control. The geological control means that the construction of the geochronological scale has been based on the dates for continuous sequences that have been most thoroughly studied and subdivided climatically and magnetostratigraphically. The chronometric control means that it is absolutely necessary to use in parallel a number of chronometric methods. The estimates obtained can be considered valid only if they have been substantiated by a n independent method for dating.

233

Taking into account these criteria of reliability, the Wiirm and Riss - Wurm short scale (and hence, all the palaeoclimatic reconstructions it permits) cannot be regarded as correct. Thus, it is already the geological control that makes us to doubt the validity of 14C final series dates obtained for the beds that are undeniably interglacial. For instance, this concerns the dates for the Karakula, Estonia, and Bolshaya Kosha, the Upper Volga, beds containing the Likhvinian flora (Velichkevich and Liivrand, 1984) or for the Karangat, Ponoi and Karginski beds with the most warm-water fauna of the Pleistocene. The palaeoclimatic situations of the Middle Wurm, assuming that the interglacial beds are of the intra-Wurm period, appear to be so utterly paradoxical that they cannot be explained from a geographical point of view. Even the first attempts of applying the chronometric control of radiocarbon data that constitute a basis of the short scale have shown that a considerable number of I4C dates, including the series dates, are not corroborated by independent methods. For instance, the sediments of the last interglacial transgression (the Neotyrrhen - Karangat - Ponoi - Cape Broughton) with 14C dates between 26 and 48 - 50 ka have been dated by the U-series technique at 60 to 120 - 140 ka. The TL dates of the continental equivalents of this transgression in the Black Sea area and Normandy also lie within an interval from 130 to 60- 80 ka. The Kalinin -Early Zyryanka tills lying under the sediments of the intra-Wurm interglacial (the Mologosheksna - Karginsk thermomers) have TL age between 100 - 110 ka and 190-210 ka. In particular, in the Monchalovo quarry near the city of Rzhev that has recently been selected as a stratotype of the intra-Wurm thermochron, A.J. Shlyukov (personal communication) has dated the Kalinin till by the TL technique at more than 210 ka. Consequently, the Kalinin till like the Warthe one corresponds to isotope stage 6 and the Mologosheksna interglacial of a stratotype locality with isotope substage 5e. By the Late Pleistocene long scale based on a set of chronometric data, the last or the Wurm glaciation is correlated with isotope stages 5d - 2 and lies between 117 and 13 ka BP. New U-series dates for the sequences of the second Boreal transgression (the Beloye Sea) in the Mesen river basin (Zaton) are within the interval from 143 t o 60-44 ka (Geochronology of the USSR . . ., 1985, p. 41). The upper layers of the Skaerumhede section (Petersen, 1984) have been dated between 47.3 to 34 ka by the I4C method. That means that in the deep-sea as well as in continental sequences of middle latitudes in Europe and North America (Chebotareva and Makarycheva, 1974; Zagwijn, 1971; Dreimanis and Goldthwait, 1973) the greatest glacial advances of the Warthe and Wiirm are actually divided by a “non-glacial” interval almost 90 ka long. The above-mentioned facts allow us to suggest a new climatochronological model of the Wurm megakryomer (Fig. 8.7) including three kryomers (stages 5d, 4 and 2 ) and two thermomers (substages 5a - c and stage 3). The stratigraphic standard of the entire Wiirm megakryomer is a unique continuous sequence of sediments of organic origin from the Grande Pile swamps in Vosges, the spore-pollen zone from 3 t o 18 (Woillard and Mook, 1981). The Wurm standard in the deep-sea sediments has been thoroughly studied in cores from the Norwegian basin, particularly cores V30-97 (Ruddiman and Mclntyre, 1981) and RC9- 181 in the eastern Mediterra-

Table 8.3. Comparison of the local climatostratigraphic units of the last glaciation

W N

P

France Deep-sea record - isotope stages

Human fossils (Mellars, 1986, Valladas e.a., 1986)

Rodents biozone

1 (Chaline, 1981)

Southern part of Russian Plain

Poland

I < Altynovo loess I

Shikhov terr. . . - I1 m 14C 7.5- 14.4. Th/U 13.9 g . Gils

-

Cro-Magnon (Lascaux) M. gregolis

- - -I

i

2nd stage

ca. 16,9

f. Dicrosronyx (Cottier) strengthening cantinentality

morsky)

I 2 Id loess

II

Gravettian (Tursac) t Y

I4C 30.9. r2 excursion

3.9

=

31 - 33.4

Enotaerka regression

Aurignarian (AW ca. 32-38

Terr. S - 6 m

, ,

, ,

,

,

"C" TM

3

Jerzmanowician tools

B f

, ' , ,

Terr. 28 - 30 m

d. Apodemus subzone (Gigny 9 - 7)

"B" KM Bogunice LOOIS

5

2

"H"

Elton regression

44-52

2.

Chatelperron TL 40-42

Terr 15-22 rn r3 excursion

Mousrierian

42 - 46

of Acheulian tradition

"C"

.$

I

[ Desna

Poznan K M

m

loen. ca 20- 24

Dofinovka I1 - 111 - "Upper Humus" of Kostyenki Streletsk tools "'C 26 - 32.7

Loess with Spitsyn ash ca. 38

Y

TL 41 - 4 3 "A" TM

50-56

Atlantic forest Denticulate M.

[

0

I

8

loess

Dofjnovka I Korman mil

-

of Kostyenki Late Moustierian (Korman IV)

'Oil Voikoi -

site

soil

I

La Quina TL 60- 68 Ferrassie TL 68 - 76

(La Quina) Humid -cold b.

%

Hurras Lake Mechetka soil

____

Upper Ciirkan r l (Shurd~Oren) excurbion KTI 91 + 1 , Th/lJ 76-81

Bug loess "Cool Moustierian" (Molodova I )

M. agreslrs

1 4 Tcrr. 48 - 52 rn Aberkun Channel

Vistula 2 - Torun? Sartowice KM TL 51 - 7 1

Loess Ketrosy w e Cool steppe

a. L. Lagtrrus

~

ypzal Denticulate Moustierian 'ombe-Grenal, 36 - 55 layers)

2-C/erhrronomys IRegourdou IV) Atlantic forest I - M oe1'onoin.w (Santeney B) Humtd - cool

c . 76-115

i

TL 97 - 101

7 Lower soil

- Ilskaya

Udai loess, KTL 66 - 78

site

Russian Plain Glaciation area

i

west Siberia

-

Mua-Khaya-

Upper Valdai - Baltic Megastade

stade Bologoe - Grudas Orsha KM

Valyek K M

Dunaevo -

Novonazirnovo T M

Kuronach-Salan TM

I4C 25 1-29.6

24.5-28.8

jlt

1 14c

-

Leyastsiems KM

31.4-34.5

Lipovka- Konoshchele KM, 31.3-33

c

Kirgilyakh I = Ulahhan-Kuel17) KM

-1

'GydaTM Mammoth body, 33 5 - 36

Bugry

~

Rogachev KV

39-42 5

l Grazdanski Praspekr - Krasnaya tiorka T M

'4c > 4 9

n ;

;

-

Lhina - Shestikhino

rtade?)

15 23 -25 8 - 10 1000 - 1200 (b) > 15 26 - 28 10- 12 1000-1100 25 Lubimovka (Ibm) (a) 1 15 22 - 24 4-6 900 - 950 (b) 15 24 - 26 6-8 400 - 750 Oskol (0s) - 10- 15 24 Sevastopol (sev) 1000 - 1400 23 first phase, (a) 15-20 26 - 28 8 900- 1 I50 21 second phase, (b) 1 15 27 - 28 10- 12 17 Yarkovka Cia) (b) 15-20 25 - 27 6-8 400 - 700 15 Bogdanovka (bgd) (a) 23 -25 2-4 700 - 800 24 - 2 6 4-6 500 - 600 (b) I1 Beregovka (brg) (a) 22 - 24 1-5 800 900 600 - 700 (b) 500 - 600 a Berezan (brs) 17- 18 - 9 - 12 600 - 800 Kryzhanovka (kr) (a) 10- 1 1 21 -22 0 + 2 7 500 - 600 (b) 22 - 23 2-4 400 - 500 6 Ilyichevski (il) 16- 17 10- 12 650 750 Shirokino (sh) (a) 8 - 10 20- 21 0 + 2 5 550 650 (b) 10- 1 1 22 - 24 1-3 __ - ~ ~ _ _ _ - _ _ - _____-_* After Shchekina (1979) up to 5- 10°C ~

-

-

-

-

~~

-

~

358

do not correlate their findings with the description of the Panagiya section made by the author, we shall try t o analyze the situation. Thus, D. quinquiramus, in the author’s opinion, could be found only in SCT 27, since it is within this SCT that Ananova recognized the first appearance of phytoplankton of the Sigmalina genus, indicating the formation of a limited (short-term) link between the Black Sea and Mediterranean Sea or the Indian Ocean which occurred about 5.5 Ma ago. Findings of Ceratolithus can be placed in 55 - 75 m interval (Fig. 11.2) where marine diatoms Actinocyclus ehrenbergi and acritarch Sigmalina are recognized as the second migration of these species. This invasion of Mediterranean species can be referred to SCT 23 and is dated as 4.7 - 4.5 Ma. Thus, this migration seems t o have occured in the middle of the Zanklean transgression and thus the Miocene/Pliocene boundary can by no means be placed higher, that is below the base of the Kamyshburun Horizon, which coincides with SCT 21 and is dated at 4.3 Ma. Thus, the Messinian salinity crisis in the Black Sea would be concurrent with SCT 28, that is with the Belbek Kryochron and the erosional break between the Pontian and the Asov. The above-mentioned iron sands can be regarded as a possible Black Sea equivalent of the Arenazzolo beds. It cannot be ruled out that the duration of the break was greater, of up to 0.5 Ma. The Middle Kimmerian Kamyshburun Horizon consists of a few layers of oolite ores interbedded with sands and varves formed during the second rewashing of tobacco-coloured ores in the beach zone. These ores are correlated with the Upper Sebastopol soil complex and Yarkov soil correlated with SCT 21, 19 and 17. As evidenced by pollen spectra the climate during the Sevastopol optima was very similar to a subtropical climate. During SCT 21 the climate was particularly warm and contrasting, with decreasing precipitation amount as compared to SCT 23 (Table 11.1). During SCT 17 the climate became even more contrasting. Yarkovka soil is bright yellow red due to highly dispersed iron oxides (Sirenko and Turlo, 1986). Mediterranean forests with subtropical taxa (Jugluns, Carya, Rhus, Nyssa, Morus and others) prevail, though pines are always present. Changes in the mammalian fauna indicate the development of savanna - steppe landscapes during SCT 21 - 17; the Moldavian fauna consists of later hipparion, antelope, gazelle, mastodon, ostrich, hyena, rhinoceros (Dicerorhinus megarhinus), camel, and a number of new taxa in the fauna of small mammals: Dolomys, Pliomys, Promimomys stehlini, Pr. constantinovae and others (Adamenko et al., 1986, p. 33). The Aydar cooling is the longest during the Moldavian - Kamyshburun Age with the development of steppe and forest steppe landscapes through the Ukraine, while the Crimean valley forests lack any thermophilic taxa. There are indications (Eberzin, 1940; Karmishina, 1975) of foraminifera Cassidulinita prima yilded by the Kamyshburun Horizon which is identical to the Akchagylian Horizon foraminifera. The Aydar Megakryochron (SCT 20- 18) seems to the author the only time appropriate to this hypothetical link between the two water basins. The climate changes occur in SCT 16 corresponding to the Pantikapeian Horizon yielding Dreissena supracimmerica and Cypriu kurluevi (Karmishina, 1975); this indicates certain water freshening. The continental section has a corresponding Kizylyar Horizon of loess-like loamy clay with steppe spore-pollen records and a new Kotlovina (Konstantinova, 1965) or Skortseli (Alekseeva, 1978, 1982) mammalian fauna with Eguus robustus, E. stenonis, Archidiskodon cf. rumanus, Mimomys polonicus and Dolomys milleri.

359

Thermo-SCT 15 (Bogdanovka) starts the Kuyalnikian - Egrissy regional stage. It is distinguished by a renewed mollusc and ostracod fauna. As indicated by Karmishina (1975) ecological recesses due to the extinction of numerous Pontian species in the Pantikapeian are substituted by Caspian Sea immigrants. The complete Early Kuyalnikian (Pokveshi, Veselovka, Skurdumi), SCT 15 - 13, over Taman and Georgia is associated with the Gauss magnethem (Fig. 11.3). The Early Skurudumi in Georgia is characterized by polydominant forests with the presence of such evergreens as Carya, Comptonia, Aralia, and so on (Shatilova, 1984). The Ukraine is covered by savanna/forest vegetation (Fig. 11.4) with broad-leaved trees in the river \alleys and isolated evergreens (Sirenko and Turlo, 1986). The Siver cooling is evidenced by a thick (2 - 8 m) layer of loess-like clay overlaying the Vaduluyvoda Terrace XI. The Gauss/Matuyama reversal can be distinguished in its upper part (Veklich et al., 1984a,b). This allows us to refer to the Siver as a megakryochron (SCT 14 and 12). The Skurdumi beds in the marine section reflect SCT 14, while the Etseri beds are referred to SCT 12. In the Early Etseri (SCT 13) dark coniferous prevail (spruce, fir, Tsuga) indicating a moderate and humid climate. Siver cooling considerably changed the composition of the Skortseli fauna giving place to the steppe Khapry one with a greater number of horses, with the first appearance of Archidiskodon grornovi, Villanyia petenyii and Mimoinys pliocaenicus. Intermediate warming in the Ukraine (SCT 13) is demonstrated by the dire/meadow soils appearing in the Siver, together with meadow/forest and brown forest/steppe soils. During the Siver the Ukraine was covered with steppe landscapes, while during the intermediate warming they were substituted by steppe/forest and coniferous forest with spruce, fir and broad-leaved trees (Sirenko and Turlo, 1985). The most dramatic event of the hydrological history of the Black Sea during the Siver Age was its direct link with the Caspian Sea on one hand and with the Mediterranean Sea on the other. Paleomagnetic studies of the Veselkovka section allow us to date these links by the time of the Gauss-Matuyama reversal. Immediately above the reversal in this section the Caspian mollusc fauna with Cardium dombra and Avimactra subcaspia substitutes for the Kuyalnikian mollusc fauna with Limnocardium (Zubakov, 1974). The beds yielding Akchagylian fauna are recognized in the Black Sea as the Polivadin Horizon. The Akchagylian fauna was discovered at the coast of the Aegean Sea by Taner in 1979 (Keraudren, 1979; Nevesskaya et al., 1986). Mediterranean plankton including Discoaster penlaradiatus, D. brouweri, Reticulofenestra pseudoumbelica, which are indicative of zones NN 17 and NN 18 appear in the Black Sea and Caspian Sea simultaneously. These species were first recognized by Lulyeva in 1979 in the core material from Chegerchi Borehole. (Semenenko and Pevzner, 1979). The location of the link between the Black and Caspian seas, however, is still uncertain. It cannot be the Manych Strait (Rodzyanko, 1984). This passage seems to be located further south, approximately along the line Elkhotovo - Sablya - Kuban. Superclimathems 11 - 9 are referred to as the Beregovka warming, when over most of the Ukraine brown forest/steppe meadow soils substituted for red savanna soils. The pollen spectra show arboreal vegetation with the prevalence of oak (14 - 27%), birch, alder, hornbeam, linden, and pine, with the addition of subtropical flora such as Carya, Celtis, Morus, Rhus, Ilex, Plerocarya. Mammalian fauna of the Beregovka Age (at least SCT 11 and SCT 10) is represented by the

360

36 I

typical Khaprovian complex. SCT 9 is associated with the Psekups fauna which is known to be transient between the Khaprovian and Odessa faunas, exhibiting the earliest Archidiskodon meridionalis (Baigusheva, 1984; Lebedeva, 1978; Rodzyanko, 1984). The Late Kuyalnikian period with the development of mollusc fauna Kuyalnikian s.str. (“Odessa Kuyalnikian”) corresponds in the marine section to the Beregovka Megathermochron in which Tretyak (1983) distinguished the Olduvai normal polarity event. Marine sediments near Kryzhanovka Village yielded bones of Hesperoloxodon antiquus cJ ausonius and a number of Dolomys millerr, D. hungaricus, Pliomys, Mimomys and others. All these were grouped by Shevshenko (1965) into the Kuyalnikian complex. It should be mentioned here that the Georgian geologists refer the Late Kuyalnikian either completely (Kitovani and Imnadze, 1974) or partially (Taktakishvili, 1984) to the Gurian. In fact, paleomagnetic studies of the sections in West Georgia in the vicinity of Meria-Gogoreti (Fig. 11.3), carried out by Kochegurova and the author (Zubakov, 1974), have shown Digressodacna digressa, D. podoliciformis fauna to have been dominant during the Gurian. I t appears at the level of the Olduvai event, which is considered to be the stratotype of the Kuyalnikian in the sections of the Kuyalnik Estuary. In any case, the marine equivalent of the Beregovka megathermochron in West Georgia is represented by three horizons, namely the Dreissena, Meria and Nadarbazevi. The first one contains the fauna Dreissenrr polymorpha weberi; Dr. rostriformis colchica and or hers (by Taktakishvili, 1984). This horizon is represented by the Tsikhisperdi beds, and it coincides with the Reunion normal polarity event and corresponds to SCT 1 1 . According to Taktakishvili (1984), this horizon tops the Egrusi, a new regional stage distinguished by Taktakishvili. The Nadarbazevi Horizon, or the lower layers of Digressodacna zone correlates with the Late Beregovka optimum (SCT 9). The intermediate cooling named Meria cooling (Zubakov, 1974) corresponds to a wellpronounced water freshening in the Black Sea. In Georgia it is recorded as varves of Khvarbeti beds with Micromelania and Pirgula, in the nearby Azov region with Tup-Dzhankoy beds with Coretus corneus and Planorbis. The Meria Kryochron is between the Reunion and Olduvai events; its calculated age is 2.0 - 1.82 Ma. Kryo-SCT 8 is associated with the formation of the Berezan loess horizon overlying flood-plain IX and of the Upper Odessa Kuyalnikian marine sediments. The stratotype is characterized by mammalian fauna with Archidiskodon meridionalis, Equus stenonis major, Elasmotherium sibiricum, Allophaiomys pliocaenicus, Fig. 1 I .3. Magnetostratigraphic correlation of the Black Sea region Pliocene sections (after Zubakob, 1974). Sections: Taman-Kerch region: 1 - Chokrak Lake; I1 - Veselovka; I11 - Panagia the Duaba river; V - Gogoreti 1 , 3 and 4; V I - Gogoreti 2: Vf1 Cape-Kutrya; West Georgia: IV - Shava and the river Chakhvata; VllI - Tsvermagal Mountain; IX - Khvarbeti 2 and 3; X - composite section. Legend: Magnetic polarity zone: 1 - normal; 2 - reversed; 3 - anomalous; Lithology: 4 - conglomerates and sands; 5 - sandstone; 6 - clay; 7 - laminated aleurites; 8 - iron ore; 9 buried soils; Fauna and flora: 10 - marine molluscs; I 1 - molluscs of brackish basins; 12 - Ostracoda; 13 floristic remnants; 14 - phytoplankton; 15 - erosional hiatuses. ~

~

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Villanyia feijari, which were studied and grouped into the Odessa complex by Shevchenko (1965). The Berezan Kryomer is reliably correlated by this faunal complex with the Lower Goryanka Horizon in the Don area and the Kobuska Horizon in Moldavia. The palynological evidence indicates that during SCT 9 the whole area of the Ukraine and the Don drainage basin was covered by wormgrass - mari steppe. In the southern Ukraine there are saltmarches with Frankeniaceae. Valley forest was mixed with limited broad-leaved taxa. On the basis of the findings of Scandinavian crystalline pebbles in the Middle Goryanka beds of the Upper Don, Krasnenkov et al. (1987) assume that in the Early Goryanka Age (SCT 8) the glacier reached the Volga - Oka area for the first time, and then melt waters brought erratics to the Don valley area. Thermo-SCT 7 (Kryzhankovka) corresponds to the formation of the alluvium Dniester terrace IX (Khadzhimus) with the Boshernitsa mollusc fauna (with

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Fig. 11.4. Paleoclimate and humidification change in the Pliocene Ukraine (after Sirenko and Turlo, 1986). I . Climatomers symbol: sh - Shirokino; il - Ilyichevsk; kr Kryzhanovka; brs - Berezan; brg Beregovka; siv - Siver; bgd - Bogdanovka; kzj - Kyzyl- Yar; ja - Yarkovka; ai - Aidar; sev Sebastopol; 0 s - Oskol; Ibm - Lyubimovka; slg - Salgir; iv - Ivankovka; blb - Belbeck. 11. The types of climate: 2 - periglacial; 3 - cold; 4 - moderate-cold; 5 - moderate; 6 moderate- warm; 7 - warm; 8 - transitional to subtropical; 9 - subtropical. 111. Precipitation in the Ukraine (mm): (a) northern zone; (b) southern zone. I V . Precipitation (rnm). V. Dominating soils in the Crirnea: 1 - reddish-cinnamonic soils of subtropical open woodland; 2 - reddish - brown soils of subtropical steppe; 3 - reddish-cinnamonic calcareous soils of subtropical open woodland; 4 - red-brown soils of subtropical steppe (in the end) and savanna; 5 - red calcareous soils of subtropical forest; 6 - red-cinnamonic soils of subtropical forest; 7 - cinnamonic meadow soils; 8 - red soils of humid subtropical forest and subtropical forest with alternating humidity. ~

363

Bogatschovia sturi) and Late Odessa (Kair) or Early-Taman mammalian fauna with Archidiskodon rneridionalis later species, Mirnomys intermedius, Prolagurus arankae, Ailophajomys piiocaenicus-laguroides and others (Lebedeva, 1972; Shevchenko, 1976; Aleksandrova et al., 1984; Shushpanov, 1983). Rich iron oxides is a distinct feature of red-brown Kryzhanovka buried soils. Pollen records give evidence of widespread forests and shrubs during the two optima of this thermochron. The forests consisted of pines, various broad-leaved species like Jug/ans, Carya, Pterocarya, Nyssa, Ostrea, Cellis, as well as dark coniferous elements. Mean winter air temperatures never dropped below zero, precipitation amounts reached 500 mm (Fig. 11.4). It is not clear as yet what is equivalent to SCT 7 in the Georgian marine section. A distinct warming stage is recognized in Chakhvata (Natanebi) Beds with dominating Didacna pavlovae and D. guriensis (Kvaliashvili, 1976). Their stratigraphic position, however, (Guriya, SCT 7 or Chauda, OCT 25), is not certain. Kryo-SCT 6 completes the climatic events in the Black Sea Pliocene. It is the time of formation of the Ilyichevsk loess horizon overlying alluvial terrace IX (Khadzhimus in Moldavia, Tanais in the Azov region) which is characterized by the Kair fauna (transient to Odessa - Taman fauna) with Ar. rneridionalis tamanensis and Prolaguruspraepannonicusprimae. SCT 6 in the Black Sea area is signified by the ingression of middle Apsheronia waters, bringing to the Azov Sea Caspian fauna with Apscheronia propinqua. This period of water freshening is associated in Georgia with the Kvemonatanebi Beds whose age is discussed. Thus it is possible to distinguish 30 climatic events in the Black Sea Pliocene, each continuing for 150 - 250 ka. Most of these events (excluding those superclimathems which are included into Sevastopol megathermomer - SCT 19-23) have distinct ecological features and can be recognized in the course of comprehensive studies of the sections. In fact, the reconstruction of climatic parameters is possible for each superclimathem; Sirenko and Turlo were the first to attempt it (Fig. 11.4). However, there is a certain difference between the qualitative estimation of the climate which allows the development of paleoclimatic time scale and their quantitative estimation. The methods for the latter one rather rough, and the estimates are only approximate. Thus, in our opinion, the estimates for larger precipitation in the Middle and particularly in Early Pliocene thermochrons (precipitation, amounting to 1000 mm and more) (according to Sirenko and Turlo) are overestimated, at least by onethird. 11.2. The Caspian Sea region

Figure 11.5 presents a composite stratigraphic scheme for the Caspian Pliocene based on a great body of published information and on data obtained by the author in Azerbaijan and the North Caucasus (Zubakov and Kochegura, 1971; Zubakov, 1973, 1974). Column 2 is mainly based on information from Rodzyanko (1981, 1984), Lebedeva (1978) and Krasnenkov et al. (1987); Column 3 is based on data of Khramov (1957), Izmail-zade and co-workers (1967), A.A. Alizade and coworkers (1972), K.A. Alizade et al. (1972), Pashaly et al. (1973), Lebedeva (1978), Trubikhin (1978), Ganzey (1984), Nevesskaya and Trubikhin (1984); column 4 is constructed from information published by Kirsanov (197 I), Karmishina (1975), Zhidovinov and co-workers (1984, 1987), Goretsky (1964) and Bludorova and co-

364

workers (1985, 1987); column 5 uses data reported by Yakhimovich and co-workers (1965, 1981, 1983, 1984, 1987). The first question to be answered is: When and why did the Caspian Sea become an isolated basin? In the Early - Middle Cenozoic the territory of the present Caspian Sea was part of the ocean Thetys, closed as the result of the collision of the Afro-Arabian, Indian, and Euro-Asian plates. The main orogenic movement occurred in the Miocene. The mountain structure of the Alpine belt formed during the Sarmatian and Maeotian. The separation of the Eastern Parathetys into two basins - the Black Sea and the Caspian Sea - took place in the middle of orthomagnethem 6, at about 6.5 Ma. Up to that time the scope of paleorelief reached, according to Sidnev (1986), 1 km in the northern Caspian Sea area. During the first few hundred thousand years of isolation (presumably during SCT 33), the residual Babajan brackish water basin still continued to exist. Since the second half of magnethem 6 a sandy productive sequence started to accumulate in the southern Caspian Sea area. In Turkmenistan the sandy sequence is named the Cheleken, or the Krasnotsvetnaya (red beds), or the Torongly formation. Its thickness ranges from 1,500 t o 3,000 m. It is characterized by a poor ostracod complex, resembling the Pontian type at the base (Bakunella, Pontoniella, Caspiolla, Xesteleberis), and the Akchagylian type at the top (Limnocythere, Leptocythere and the like). Therefore Kovalevsky (1936), Popov (1967) and Karmishina (1975) do not regard the productive sequence as a separate stratigraphic unit. A change in the type of ostracod fauna is associated with the middle part of the Balakhany formation, i.e. with the boundary between polarity zones 4 and 5 in Izmail-zades (1 967) interpretation, or with the lower Gilbert polarity reversal, at 5.38 Ma, in the author’s opinion. The Pontian part of the productive sequence is divided by the Azerbaijan geologists into four climato-sedimentary units (Fig. 1 1.5). The hydrologic regime of the Kala - Kirmakina basin is a subject of controversy. The presence in its sediments of gypsum, anhydrite and celestine suggests high salinity (Alizade et al., 1972), while the ostracod fauna gives evidence for freshwater conditions. Kvasov (1966) believes that the Kirmakina basin was not merely a salt water, but a supersalt water basin, and freshwater fauna inhabited not the basin, but its limans and lagoons, into which rivers discharged. This reconstruction appears to be more sound. The point is that the base of alluvium of the great Kinel River, a precursor of the Volga and Kama rivers, which fell into the basin of the productive sequence, had been 540 m, 100 m and 2- 10 m below sea level at Astrakhan, Kazan and in the Kama - Vychegda lnterfluve, respectively (Sidnev, 1986). Hence, the basin itself must have been no less than 600 m below sea level and could not but be an evaporite basin. However, freshwater basins may also have existed, provided a sufficiently intensive freshwater run-off occurred. Not one basin but a system of semi-closed basins with differing hydrological regimes may have existed in the Caspian depression during Late Pontian time. Here the evolution from the Pontian fauna to the Eoakchagylian may have taken place. Narrow erosional cuts of the Kinel River and its tributaries are filled in with alluvium having, like the Kirmakina formation, normal polarity (Fig. 11.6). Pollen analysis and the finding of fossil antelope in the Chebenki I horizon suggest the existence of steppe landscapes in the territory of

P

Isl

Chrbenki

Fig. 11.5. Stratigraphy and climatic events sequence i n the Pliocene Caspian region. 1 - magnetostratigraphic scale (after Berggren et al., 1985; Mc Dougall et al., 1984) with the author's corrections; 2 - the Don valley (on the left) and Azov - Manych -Terek plain (after Krasnenkov et al., marine section of Transcaucasia (after the author; 1987; Lebedeva, 1978; Rodzyanko, 1984); 3 Alizade K.A. et al., 1972; Izmailzade et al., 1967; Lebedeva, 1978); 4 - Northern Caspian region (on the left) (after Zhidovinov et al., 1984, 1987) and Middle Volga and Lower Kama buried valleys (on the Pre-Urals region (after Yakhimovich et al., right) (after Goretsky, 1964; Bludorova et al., 1987); 5 1981, 1987). ~

~

Bashkiria (Yakhimovich et al., 1965, 1977). According to Ananova (1974), in the lower reaches of the Kama River steppe gave way to dark coniferous taiga with minor hemlock, keteleria, nissa, liquid amber, and other species, which could survive neither frost nor drought. O n the Apsheron Peninsula the Pontian and the Akchagylian parts of the productive sequence are separated by the so-called "Pereryv" (break) formation, made up of coarse inequigranular continental sand with pebbles. It is close t o SCT 28 in age.

366

In the northern Caspian Sea area, equivalents of the Eoakchagylian part of the productive strata, i.e. those of the Balakhany formation, are: the Furmanov Clay, up to 300 m thick, according t o Kirsanov (1971); and the Verkhne (upper) Kushum formation, according to Zhidovinov and co-workers (1 984, 1987). They are characterized by a new (Akchagylian) ostracod assemblage, containing Prolimnocythere scharapovae, Citherissa, Cytherida torosa and others, a fresh -brackish water fauna of molluscs, including Dreissena-Clessiniola, as well as euryhaline marine foraminifera Bolivina, Cassidulina, Ammonia beccarii, and the like (Karmishina, 1975). In the buried valley systems of the Kine1 River, the Early Akchagylian corresponds to lacustrine - lagoonal deposits, namely to the Chelna and Chebenki 111 horizon. According to the paleomagnetic scale (Fig. 11.6) they are 5.3 - 4.5 (4.0) Ma old. The formation of lagoonal facies suggests a raising of the sea level in the Balakhany basin. Pollen statistics, reported by Ananova (1 974) for the Chelny horizon and by Yakhimovich and co-workers (1965) for Chebenki 111, suggest that at that time the Kama River basin and the cis-Ural area were covered by dark coniferous forests, containing Tsuga diversifolia and Picea minor, with minor subtropical species, including nissa, sequoia, beech, nut, and the like. Miocene relicts Salvinia glabra, Selaginella bashkirica, Morus tanaitica, Caldesia cylindrica, Euryale nodulasa and others were found by Dorofeev (1984) in brown coal measures at the base of Chelny 111. Two temperature maxima occurring during Chebenki 111 time (Fig. 11.6) may be assigned to SCT 27 and 23 of the paleomagnetic scale. The top of Chelny, coinciding with the Nunivak event, corresponds to SCT 21. So far, equivalents of the Chelny and Chebenki I11 have been found neither in Azerbaijan nor in Turkmenistan. But, as stated above, the Eoakchagylian fauna of foraminifera was discovered in the Black Sea basin in the lowermost Kimmerian (Eberzin, 1940). This implies that the Eoakchagylian transgression is synchronous to the infux of Mediterranean water into the Black Sea basin. Some part of the Mediterranean fauna, for example foraminifera, nannoplankton, as well as the most euryhaline forms of molluscs of the genera Cardium, Clessiniolla and others, got from the Black Sea into the Caspian Sea through a passage in the North Caucasus no later than 5.3 Ma ago, i.e. during SCT 27. Not one but several phases of Eocaspian transgression may have occurred. They coincide with the onset, middle and close of the Karlaman. The point is that nannoplankton of zone N N 15, containing Reticulofenestra pseudoumbelica, found by Lyulieva in the Middle Akchagylian of the Yasamal Valley, Apsheron, and at Mount Duzdag (Semenenko and Lyulieva, 1982), could penetrate into the Caspian Basin before it became extinct in the ocean, i.e. before 3.5 Ma; at the same time Fig. 11.6. Magnetostratigraphic correlation of Pliocene Caspian basin sections (after Zubakov, 1974). Sections: 1 - Kvabebi; 11 - Kergez; 111 - Yasamal Valley; IV - “Productive Formation” section (after lzmailzade et al., 1967); V - Kobi Valley and Degdovskikh Mountain; VI - Zykh Lake; VII Chernorechie; V l l l - Elkhotovo; 1X - the rivers Podkumok and Baksan; X - composite section (after the author). Magnetic polarity zone: 1 - normal; 2 - reversed; 3 anomalous. Lithology: 4 - conglomerates and sands; 5 - aleurites; 6 - clays; 7 - till-like loams; 8 - lavas and ignimbritic bombs; 9 - ashes. Fauna: 10 - shell-limestone; 11 - freshwater molluscs; 12 - Ostracoda; 1 3 - mammalian remnants; 14 - carpologic fossil flora; 15 - hiatuses. ~

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Kuyalnikian molluscs, such as Amphimelania impressa, found by Goretsky (1964) and Yakhimovich and co-workers (1965) in the Sokol horizon and in the Karlaman, could not enter the Caspian Sea before the onset of the Kuyalnitian, at 3.1 Ma, or, at least, before the beginning of the Pantikapeian, at 3.4 Ma. Passages between the basins probably recurred during SCT 22, 18 and 16. Are not alternations in the Black Sea section of Kimmerian ore beds, containing the Ponto-Dacian fauna, with clay beds, having almost freswater fauna, including Akchagylian fauna, the best evidence for the existence of these passages? Paleontological data on the lower Sokol horizon (Ananova, 1974; Bludorova et al., 1987) and on the Karlaman (Yakhimovich et al., 1965, 1981) show that the transgressive phases of development of the Early Akchagylian basin coincide with coolings. This is suggested by the taiga character of pollen spectra and the appearance at the top of the Karlaman of boreal foraminifera Cribroelphidium heterocameratum Volosh. and Elphidium subarcticum Cusch. (Yakhimovich et al., 1965), as well as grass Zostera nana Roth, inhabiting sea water (Dorofeev, 1984), which could penetrate into the Kinel River Basin, according to Yakhimovich and coworkers (1981) only from the north, from the Pechora River basin. It cannot be ruled out that the Kolva transgression in the Pechora basin is indeed Middle Pliocene in age (Zarkhidze, 1981; Baranovskaya and Zarkhidze, 1985) and that the Polar basin and the Caspian basin may once have been linked directly. If the basins had really been linked in the past, the linkage is likely to have taken place during SCT 16. On the Apsheron Peninsula the base of the Surakhany formation is equivalent to the Lower Akchagylian transgression. The top of the Surakhany formation, from the Kaena episode to the Gauss/Matuyama polarity reversal, characterized by an especially shallow water ostracod complex with C . littoralis (Zubakov and Kochegura, 1971), is recognized as the Kergez regression, equivalent to the Middle Akchagylian in the sense of A.A. Alizade and co-workers (1972) and to the Kumurla, according to Yakhimovich and co-workers (1965). The Kergez regression consists of two phases, probably coeval t o SCT 15 and 13. The Eruslan ingression, marked in the northern Caspian Sea basin, started just after the Kaena excursion (Zhidovinov et al., 1984). The Kumurla is best characterized by pollen spectra from the Borehole 40 section, near the settlement of Rybnaya Sloboda, Kama River, where this time interval is assigned to the upper part of the Sokol horizon. In the lower section, preceding the Mammoth excursion, there was recognized a pine - fir pollen complex, dominated by Pinus sect. Eupitys and P. sect. Eupicea and two hemlock species: Tsuga canadiensis and T. diversifolia; broad-leaved trees, including hornbeam, beech, oak, elm and lime, accound for 5% of the complex (Bludorova et al., 1987). This thermochron is synchronous to SCT 15. The second pollen complex, characterized by the drastic predominance of fir (up to 42%) and the disappearance of most broadleaved trees, except for lime, corresponds to SCT 14. Yakhimovich and co-workers (1965, 1981, 1984) believe that the first appearance of mountain tundrain the Urals, as suggested by pollen of Lycopodium pungens, L . appressum, is associated with intra-Kumurla cooling. The third pollen complex from the Rybnaya Sloboda area, whose spectrum suggests the presence of fir and broad-leaved trees, including oak, elm, maple and hornbeam, is associated with SCT 13.

369

In the Volga River basin “warm” mixed coniferous and broad-leaved spectra with hemlock were recognized in the lower part of the Eruslan (SCT 15) and Urda (SCT 13) transgression cycles, and “cold” steppe spectra were ascertained in the upper part of these cycles. Flora and fauna of the Simbugino section, Bashkiria, and, in particular, the first appearance in the Caspian Sea area of lemming (Synoptomys mimomiformis) imply cooling, and the presence of Mimomys pliocaenicus indicates the Khaprian age of the fauna. At the same time all organic remains were taken from normally magnetized beds (Yakhimovich et al., 1987). This somewhat confuses the dating of the Simbugino kryochron. Possibly, it could be assigned to SCT 14 (or the SCT 13/12 boundary). Mammalian fauna from the Kvabebi section, East Georgia (Gabuniya and Vekua, 1968) derives from normally magnetized (Fig. 11.7) Middle Akchagylian sands above ash with fission-track age at 2.55 ? 0.20 Ma (Ganzey, 1984) to 3.0 Ma (Ushko et al., 1987). It includes Kvabebihyrax kacheticus (daman), Anancus arvernensis, Dicerorhinus megarhinus, Hipparion ex gr. crussafonti, Parastrepsiceros socolovi, Ioribos aceros. In the experts’ opinion, this fauna is similar to the Etouarian fauna in age, and its ecological composition corresponds to a landscape of moist savanna. Now the Kvabebi area is situated within the desert. The Kvabebi

rr

1

Fig. 11.7. Pliocene climatic events in the Volga-Urals region (after Yakhimovich er al., 1984). 1 - 2 regional stages and horizons; 3 - paleomagnetic scale; 4 - section division in the Northern Caspian region and Lower Volga region (after Zhidovinov et al.) and vegetation change, in the Lower Volga section division in the Middle Volga region and Kama region (after region (after Chiguryaeva); 5 Bludorova et al.) and vegetation changes (after Bludorova); 6 - section division in Pre-Urah (after Yakhimovich et al.) and vegetation changes (after Nemkova); 7 - Caspian transgressions (after Yakhimovich); 8 - super- and hyperclimathems (after the author). -

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fauna, as well as similar fauna from the lower horizon of the Kushkuna section, Azerbaijan, which contains Anancus arvernensis, gazelle, deer, tortoise, also collected from regressive Middle Akchagylian sands (Lebedeva, 1978), cannot be dated more precisely than as the upper part of the Gauss epoch (Zubakov, 1974). It seems wise to consider it as the undifferentiated Kvabebi mega-thermochron (SCT 15 to 13). To its close is assigned alluvium of the Kyzburun sequence, which in the Baksan River basin separates ignimbrite with K/Ar ages of 2.4 and 3.0 Ma and the Baksanges sedimentary-volcanic sequence with a fission-track age of 2.2 Ma (Zubakov, 1974). The maximal development of the Akchagylian transgression (Middle Akchagylian in the sense of A.A. Alizade and co-workers, 1971, and Upper Akchagylian in the sense of Nevesskaya and Trubikhin, 1984) is synchronous with the lower part of the Matuyama magnethem (Fig. 11.6) and SCT 12. Nannoplankton, including Discoaster brouweri and D. pentaradiatus, i.e. species, characteristic of zones NN 17 and 18, dated at 2.65 - 1.88 Ma, has recently been found in the beds (Semenenko and Lyulieva, 1982). A tooth of the Khaprian elephant, Ar. gromovi, was found by Lebedeva (1978) in the Kushkuna section, in deposits transitional from the maximal phase of transgression (SCT 12) to the regressive, Upper Akchagylian, phase (SCT 11). Ash just below the finding is dated by the fission-track method at 2.19 k 0,18 Ma (Ganzey, 1984). This is in good agreement with the age of the SCT 12/11 boundary, estimated at 2.23 Ma. As stated above, water of maximal transgression, which passed into the Sea of Azov through the inferred Elkhotovo - Kuban passage, was dated on the Taman Peninsula by the GausdMatuyama reversal - it is somewhat younger than this reversal (Zubakov, 1974). A slightly different picture is inferred from the data available on the Volga River basin and cis-Ural area. Here the Chistopol and Zilim-Vasiliev horizons, coeval to SCT 12, are regarded only as the initial part of the transgressive cycle. The maximum of the transgression, marked by the furthest advance of marine beds along the Kine1 River valleys, is associated with warming dated by the Reunion normal polarity excursion (Yakhimovich et al., 1987). But on the Taman Peninsula a regressive sandy unit (Fig. 1 1.3) with Avimactra subcaspia and Kuyalnikian rodent fauna is fixed at the Reunion level (Zubakov, 1974). In the author’s opinion, controversy could be avoided by correlating the Akkulaevo normal polarity episode not with the Reunion (2.13 Ma), but with the Argentine excursion (2.23 or 2.33 Ma). Palynological data suggest that the climate of the Chistopol - Zilim-Vasiliev kryochron corresponded to the condition of the present taiga zone characterized by a rhythmic change of a dry temperate cold climate (pollen of pine, birch, alder) by a wetter and warmer climate, when taiga with lime and hemlock existed there (Ananova, 1974; Bludorova et al., 1987). SCT 12, as well as subsequent SCT 11 and SCT 10, are characterized by the Akkulaevo ( = Khaprian) mammalian fauna with Archidiskodon gromoviand Mimomyspliocaenicus (Yakhimovich et al., 1983). The Simbugino fauna of freshwater molluscs with Potomida baschcirica and the like, described by Chelpalyga (1980), is also coeval with SCT 12. In the Elkhotovo section, where the Terek River descends the mountains and flows across the plain, the marine Akchagylian is interbedded with till-like mudflow (Fig. 11.7), previously taken for till. Mudflows, extending for a few tens of kilometers within the valley, suggest alpine glaciation of the Caucasus. Milanovsky

37 I

and Koronovsky (1969) marked more intense volcanic activity in the Greater Caucasus range during Middle Akchagylian time. Radiometric and paleomagnetic data allow the association of the volcanic maximum with the time interval 3.2 to 2.2 Ma (Zubakov, 1974) and synchronization of ignimbrite “explosions” with phases of alpine glaciation of the Caucasus and with Caspian transgressions. For example, according to A.N. Komarov, the age of obsidian bombs of the Baksanges formation, estimated at 2.2 k 0.20 Ma (Zubakov, 1973), i.e. close to the age of Kushkuna ash, 2.19 +- 0.20 Ma (Ganzei, 1984), unambiguously indicates a temporal relationship with SCT 12. In the Volga - cis-Ural area thermo-SCT 11 is reliably divided into three paleoclimatic stages. Thermochron 1 l c is simultaneous with the Late Akkulaevo Early Uzen retreat of the sea and with the existence of coniferous and broad-leaved forest and forest steppe. The thermochron is associated with the Sultanaevo faunistic assemblage of the Levantine-type freshwater molluscs, containing Bogatschevia tamanica, Rugunio samarica and Yugoslavian - Mediterranean elements, for example, Unio metochiensis and the like (Yakhimovich et al., 1983). Kryochron 1 l b is related to the upper part of Uzen beds, characterized by a steppe pollen spectra and coeval to advance of the sea and salinity decrease (Zhidovinov et al., 1984, 1987). Thermochron 11, corresponds to the Aralsor regression and to the lower part of Uzen horizon. At that time coniferous and broad-leaved forest, containing hemlock, again appeared in the Volga basin (Zhidovinov et al., 1987). Thermochron 11, is probably associated with the Early Voevodinskoe warming (Yakhimovich et al., 1983) fixed in alluvial deposits. Kryo-SCT 10 represents the termination of the Akchagylian history of the Caspian Sea basin, when the last, Late Akchagylian, according to Y.P. Kolesnikov, transgression, more freshwater as compared to the previous one, developed against the background of retreat of the sea. In Azerbaijan the transgression is associated with the so-called Geran beds, made up of lagoonal black varved mud with Micromellania, resembling the Meria beds with Micromellania and Pirgula in the Black Sea basin. In the Volga River basin and in the cis-Ural area they are recognized as the Biklyan horizon by Goretsky (1964) and the Voevodinskoe horizon by Yakhimovich and co-workers (1983). As a whole, SCT 10 in the Caspian basin, as in the Black Sea basin, was a time of substantial salinity decrease. Steppe pollen spectra of the Biklyan - Voevodinskoe horizon suggest aridization and cooling. Mammalian fauna is represented by the late version of Khaprian complex with Mimomys intermedius. Bludorova and Filicheva (1985) propose to relate the Omar diatomite also to the Biklyan kryomer. With regard to paleoclimates, the Apsheronian of the Caspian Sea basin embraces five superclimathems, SCT 9 to SCT 5. The Akchagylian/Apsheronian boundary cannot be reliably drawn on lithological and faunistic grounds; in the complexly studied sections it is drawn within the double Olduvai polarity zone (Menner et al., 1972; Zubakov, 1973, 1974; Trubikhin, 1978). On rhythmo-stratigraphic grounds the Apsheronian regional stage is divided into three climato-sedimentary cycles. Recently they have been named after local geographical names (Fig. 11.5). The lower part of all three cycles corresponds t o regressive phases of the basin and to thermo-superclimathems, and the upper part is associated with transgressions and kryo-superclimathems.

372

The Lower Apsheronian climatic cycle in sea sections (Yasamal - Novokazanka horizon) is characterized by poor fauna of a highly freshened basin with the predominance of gastropods. In the near-shore facies of Azerbaijan this cycle is associated with remains of the earlier form of southern elephant - Archidiskodon meridionalis meridionalis (Lebedeva, 1978). The lower part of the Dema alluvial formation, containing Odessa fauna and Bogatschovia sturi, is related to SCT 9 (Yakhimovich et al., 1983). In the Elkhotovo section (Fig. 11.7), in lacustrine marls, at the top of the Olduvai episode (Minaret beds) V.I. Vasiliev found leaf imprints of Pistacia miochinersis, Platanus aceroides, ferns Dryopteris cf. meyeri and the like (all-in-all 14 species), suggesting a warmer and more humid climate than that of today (Zubakov and Kochegura, 1971). The Middle Apsheronian cycle is recognized in the Kobi - Tsubuk sea sections from the acme of Apsheronian mollusc fauna - the appearance of multiple cardiides (Apscheronia propinqua, Parapscheronia raricostata, Pseudocatillus, Monodacnu and the like). A normal polarity episode, the lower one of three events in the upper part of Matuyama, is recognized at its top (Fig. 11.7). Previously the episode was correlated with the Jaramillo event (Zubakov and Kochegura, 1971); at present it is correlated with the Cobb-Mountain event, 1.1 Ma. Lebedeva (1978) relates the find of the late form of southern elephant (Ar. rneridionalis tamanensis), characteristic of the Nogaisk - Kair assemblage of the Black Sea area, to the Middle Apsheronian of Mount Duzdag (SCT 6). The lower part of the Tsubuk horizon in the Volga basin (SCT 7) yields a forest steppe spectrum with broad-leaved elements (Zhidovinov et al., 1987). The most complete characteristics of SCT 7 was obtained from the lower part of Dovlekanovo horizon, cis-Ural area. The Yulushevo fauna of freshwater molluscs with Bogatschevia scuturn and Pseudosturia caudata was ascertained here (Chepalyga, 1980; Yakhimovich et al., 1983). In the Kama River Valley SCT 7 is synchronous with the Ik horizon, containing Viviparus kagarliticus and a complex of carpoids with Trapa and Euryale europaea ( = Pseudoeuryale), resembling Tegelenian E. timburgensis (Goretsky, 1964; Bludorova et al., 1987). Pollen spectra suggesting the existence of light fir forest with minor dwarf birch were obtained by Bludorova and co-workers (1987) for the upper part of the Laishevo and Gorki formations (SCT 8 and 6) in the Kama River basin. In the Elkhotovo section, at the top of the RukhsDzuar series, composed of tuffaceous-sedimentary rocks, a mudflow horizon is associated with a normal polarity event, thus suggesting the existence of glaciers during SCT 6 in the upper reaches of the Terek River, in the Mount Kazbek area (Zubakov and Kochegura, 1971). The above review implies the same succession of climatic events in the Caspian and the Black Sea basin during Pliocene time. Moreover, the hydrological history of both basins appears t o have been closely inter-related. So far only two reliable passages have been known to provide a link between the basins in the Pliocene (Zhizhchenko, 1969). In the Middle Akchagylian (SCT 12), at 2.4 - 2.2 Ma, the flux of Caspian water, marked by molluscs Cardium dombra and Avimactra subcaspia, penetrated via the Elkhotovo - Kuban passage into the Sea of Azov and Sivash, left some fauna in sections of Romania, and reached the Aegean Sea. At that time the Caspian sea level was roughly 100 m above the present datum, and the Great Akchagylian sea more than twice exceeded the present sea in size. During the Middle

373

Apsheronian (SCT 6), at 1.2 - 1.1 Ma, water, containing Apscheronia propinqua, again flooded the Sea of AZOV,this time through the Manych Strait. We can state that both transgressions were associated with phases of the “saltwater Caspian Sea”, i.e. with the phases of its faunistic acme. Just because of this, they are clearly marked in sections of the Black Sea basin. Other influxes of Caspian water into the Black Sea, inferred by the author, may have occurred in phases of strong water freshening. Hence, they could hardly be fixed in the Black Sea sections. The existence of passages between the two basins is confirmed by findings of: Akchagylian foraminifera in the Kimmerian sections; Mediterranean nannoplankton in the Akchagylian sections; Kuyalnitskian molluscs in the Kinel River estuaries. From the above review it is inferred that a link may have been provided between the basins during kryochron stages. Against this inferrence is the relation between the peak of the Great Akchagylian transgression and the Akkulaevo horizon (Yakhimovich et al., 1965, 1983, 1984, 1987), assigned by the author to thermo-SCT 11. However, the point should be elucidated. First, the Akkulaevo excursion is still to be dated (it can prove to be older than the Reunion); second, the Akkulaevo excursion may have been coincident while a transitional, intra-SCT 1 I , cooling, equivalent to the Tegelenian “B” kryophase (Zagwijn, 1974). Thus, this inferrence appears to be valid as well. Caspian transgressions in the Pliocene are synchronous with coolings, or, to be more precise, with their first half. This is an empirical fact. And maximal transgressions (SCT 12 and 6) coincide with the greatest known Caucasus glaciations. Connections of the basins during SCT 27 - 26, 22 and 18 - 16 also coincide with global surges of coolings (see below). Caspian regressions in the Pliocene coincide with thermochrons. The lowest Caspian sea level was associated with the Pliocene temperature optima - SCT 31, 29, 27, 23 (Chebenki I, 11, IIIa, IIIc). The Akchagylian - Apsheronian regression is synchronous with the last temperature optimum (SCT 9). However, within the Caspian region proper the climate appears not to have been dry and arid in regressive phases. For example, during the Kvabebi mega-thermochron (SCT 15- 13) savanna landscapes existed in place of the present desert. During all the Pliocene thermochrons in the northern Caspian Sea area the territories of present steppe and forest steppe (in Bashkiria) were occupied by dark coniferous and broad-leaved forest with hemlock, beech, and even single taxodium. And, vice versa, during transgressions of the Pliocene Caspian Sea, its coasts were covered by steppe, while in the Kama River basin dark coniferous forest gave way to pine forest. This inference is in agreement with data of Chiguryaeva (Yakhimovich et al. 1984) on the Apsheronian. Hence, the following generalizations have emerged from the review: (1) In the Pliocene, as in the Pleistocene, the Caspian and Black Sea level variations were out of phase; (2) In the Pliocene, the Caspian sea level was very responsive to changes in evaporation; (3) Caspian sea level variations were cophasal with glaciations and, hence, with general changes in atmospheric circulation. As for the relationship between the Pliocene Caspian sea level and freshwater run-off, we can find no direct evidence which substantiates it geologically. However, indirect evidence suggests that during thermochrons (and regressions) freshwater run-off had low seasonal variations and increased in summer time. During kryochrons (and Caspian Sea transgressions) the seasonal amplitude of run-off increased due to spring floodings.

3 74

11.3. The Mediterranean, North-Western Europe and other regions In addition to the Pontian - Caspian stratigraphic standard for the Pliocene, there are about a dozen other standards, such as the Mediterranean, North Sea and Pannonian in Europe; Kazakhstan - West Siberian, Central Asian, East Siberian and North-Eastern in the USSR; Pacific, Great Plains and Atlantic in North America; Japanese, Sivalik in Asia; New Zealand and Argentine in the Southern Hemisphere. The first two schemes are undoubtedly the most elaborate. Space does not permit a detailed summary of all the standards and emphasis will be placed on correlation and discussion, proceeding from the assumption that the Englishspeaking readers are acquainted with original data. Figure 11.8 collates stratigraphic subdivisions for deep-sea sequences of the ocean (column 2 - 3), and the Mediterranean (column 4), for terrestrial sequences of south-western (column 5) and northwestern (column 6 ) Europe and attempts to tie them to the succession of superclimathems, developed for the Pontian - Caspian sequences. However, prior to discussing the proposed correlation, let us dwell on three debatable questions of the chronology of the Pliocene: (1) When did the Messinian crisis of salinity take place (its onset and end)? (2) When did the transition from the “warm” Early Pliocene (Zanclean - Ruscinian - Brunssumian) to the “temperate” Pliocene (Piacenzian - Villafranchian - Reuverian) occur? (3) What is the age of the “cold” Late Pliocene, marked by the appearance in the Mediterranean of first “northern guests”? The answers to the first two questions are especially important for paleoclimatic reconstructions, as the top of Messinian and the base of Villafranchian mark the time interval of the thermal optimum of the Pliocene sensu lata, which could serve, according to Budyko (1980), as an paleoanalogue of the future climate. Classical works by Cita, Ryan, Hsii, Berggren, Sierro and others (Berggren and Couvering, 1974; Cita et al., 1973; Hsii et al., 1977; Cita and Colombo, 1979; Ryan et al., 1973, 1974; Chumakov, 1974, 1982; Sierro et al., 1987) led to the conclusion that the Messinian crisis of salinity occurred in the time interval from 6.5-6.6 (6.6-6.2) to 5.4-5.3 Ma. The universally known biomarkers of the Early Messinian and the Late Messinian are the appearance in the Mediterranean Sea of COCcolith Amaurolithus primus and colder water foraminifera Globorotalia conomiozea, respectively; the close of the Messinian is marked by a drastic change of the Lago Mare euxinic biota, containing Pontian - Pannonian Congeria and Cyprideis, by Atlantic sea biota, including Sphaeroidinellopsis sp. and Globorotalia margaritae. The main data supporting this version were obtained from Andalusia, Morocco and DSDP Hole 397, since in the type region on Sicily the sections are unsuitable for paleomagnetic dating. Recently a group of scientists, who studied the sequences on Crete (Langereis et al., 1984) and then in Calabria and on Sicily (Hilgen, 1987), worked out a new version reducing the age of the Messinian event and narrowing the time interval to 5.6 - 4.85 Ma. The revision was based on new correlation of findings of GI. conomiozea on Crete with the paleomagnetic scale.

375

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Fig. l l .8. European Pleisticene chronology - synoptical table (author's version). 1 magnetostratigraphic scale (Ma) (after Berggren et al., 1985; McDougall et al., 1984) with age corrections; 2 - zonal subtropical ocean waters division according to calcareous foraminifera (after Blow, 1969) with boundary age corrections; 3 - the same according to nannoplankton (after Martini, 1971) with age corrections; 4 - composite calcareous plankton scale for the Mediterranean (after Rio, Sprovieri and Raffi, 1984, with additions after Driever, 1984; Hsii et al., 1977). numbers 1 - 8 mark integrated calcareous - plankton intervals (ICPI), combining foraminifera and plankton zones; 5 - age subdivision according to mammalian fauna (after Azzaroli, 1977, 1983; Berggren et al., 1985b; Demarcq et al., 1983; Shaline and Michaux, 1977; Bonadonna and Alberdi, 1987; Steininger et al., 1987); 6 North-eastern Europe (the Netherlands, Great Britain, FRG) according to the data of Van der Hammen, Wijmstra and Zagwijn (1971), Zagwijn (1974), Menke (1975), Meyer (1981). Boenigk et al. (1974) and others; 7 - superclimathems and hyperclimathems, their calculated ages (Ma). ~

316

Langereis and co-workers (1984) believe that this Antarctic immigrant, whose evolutionary appearance in DSDP Hole 284 is dated at 6.15 Ma (Loutit and Kennett, 1979), appeared in the Mediterranean only half a million years later during the reserved polarity event of magnethem 5, i.e. at 5.6 Ma. Although new detailed isotope magnetic data on deep-sea sediments of the ocean are already correlated with the version of Langereis and Zijderveld et al., this version seems extremely vulnerable of criticism. The reason is not only the absence of radiometric substantiation of the new version, but the arguments used by Langereis and co-workers (1984). The point is that the evolutionary replacement of Globorotalia miozea by GI. conomiozea, which occurred near the SCT 31/30 boundary (Fig. 11.9), is no doubt associated with the strongest surge of Cenozoic cooling, resulting, particularly, in the formation of ice sheet in the Patagonian Andes about

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Fig. 11.9. Paleoclimatic events o n the Miocene - Pliocene boundary in the Southern Hemisphere high latitudes, recorded in a section o n the river Blind, South Island, New Zealand (after Loutit and Kennett, 1979, fig. 1). 1 - New Zealand stages; 2 - depth(m); 3 - sample numbers; 4 - magnetic polarity (black parts mark reversed magnetic polarity); 5 - orthomagnethems; 6 - age (Ma); 7 - stable isotopes; 8 relative sea level; 9 - bioclimatic indicators (1 - Neog/oboquadrina pachyderma sinistral factor - cold; 2 - G/obigerina bulloides factor - warm; 3 - frequency in 0711 N. pachydermu - cold); 10 - intervals of reference species spreading. Temperature peaks, numbered from 25 up to 32 and their correlation with superclimathems (after the author).

311

6.5 -6.1 Ma BP (Mercer and Sutter, 1982). The first “cold guest” could penetrate into the Mediterranean only owing t o this surge of cooling, i.e. before the close of SCT 30, whose upper boundary coincides with the magnethem 6 / 5 paleomagnetic reversal. Figuratively speaking, this guest could hardly have reasons to delay its expansion for 400 ka, to the next surge of cooling, i.e. to SCT 28. From the bioclimatic point of view, GI. conomiozea “was” to appear in the Mediterranean no later than 5.95 Ma, i.e. at the end of magnethem 6, as accepted by Cita and other authors. Therefore the author agrees with Berggren (1987), who believes that the FAD age of GI. conomiozea in the Mediterranean, estimated by Langereis and coworkers (1984) at 5.6 Ma, is insufficiently substantiated. It should be noted, also, that correlation of the Mediterranean events with the Black Sea and the Caspian Sea events gives evidence for the traditional age estimate of the base of the Messinian at 6.5 - 6.0 Ma. First, the Pontian and Lower Messinian deposits are very similar in lithology and biostratigraphy. Clauzon (1981) recalls that the Cucuron marl in the Durance basin and the overlying beds with Congeria ( = Lago Mare) had been once described by Deperet as stratotypical for the Pontian. Clauzon (1 981) thinks that mammalian fauna with Hipparion, occurring in these deposits, as well as the Taurian fauna from the Pontian deposits of the Pontian - Caspian basin, belong to zones N N 12 and 13. Second, the most dramatic events in all three basins - their “drying up” - fall within kryo-SCT 28, at 5.7 - 5.5 Ma. In Morocco their equivalent is the Late Messinian Bou Reg Reg ocean regression, dated at 5.6 Ma (Cita and Colombo, 1979). In the light of these data, the fact that the Kimmerian and Eoakchagylian transgressions wert synchronous with the Zanclean transgression becomes almost evident. True, it should oe said in all fairness that the presence in the Black Sea section of the Kutrya lacustrine beds within the interval from 5.7 to 4.7 Ma still leaves a chance for further discussion of the version of Langereis and co-workers (1984), as applied to the Black Sea. The version of Langereis and co-workers should not be identified with the concept, worked out by Chepalyga (Regional Committee. . . 1985, p. 139), about the association of the Messian crisis in the Black Sea with a hypothetical Pontian/Maeotian break. The hypothesis appears to be unsubstantiated and hence there is no point in discussing it. Let us direct our attention to correlations of the Messinian crisis with oceanological events, recently elaborated by Hodell and co-workers (1987). In the succession of events, proposed by the above authors (Fig. ll.lO), the age of the Messinian is accepted from the version of Langereis and co-workers, i.e. at 5.6 - 4.9 Ma. This correlation is open to criticism and ambigous, as seen from triangles denoting the position of the Pliocene/Pleistocene boundary, accepted in each borehole. It corresponds to the Mediterranean - Black Sea standard only in DSDP Holes 397 and 516 A and, to a lesser degree, in DSDP Hole 590. But that is not the point. The figure is presented here to show that the number of main climatic events (superclimathems), fixed in six boreholes, drilled in two oceans, within the time interval from 6.3 to 4.5 Ma (to the Sidufjall event) is the same as in the above three regions. The second debatable question concerns the “warm”/“temperate”, i.e. the Lower/Middle Pliocene boundary, in marine and terrestrial sequences. Two opin-

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Fig. I 1 .lo. Oxygen-isotope curves of benthic foraminifers from sites 588 (south-west Pacific), 397 (north Atlantic) 590, 284 (south-west Pacific), 516A and CHI 15 (south-west Atlantic). The Miocene- Pliocene boundary, defined by different methods, is indicated by a triangle (after Hodell, Elmstrom and Kennett, 1987, fig. 2).

3 79

ions exist on this question. For brevity, we discuss them together. Some scientists hold to the value of approximately 3.3 - 3.2 Ma for the ZancleadPiacenzian boundary, coinciding with the disappearance of thermophilic Globorotalia margarifae and the appearance of temperate water dweller GI. crassaformis; the Ruscinian/Villafranchian boundary, coinciding with the appearance of developed fauna of Microtinae with Mimomys polonicus and horse; the Brunssunian/Reuverian boudary, coinciding with the disappearance of taxodium forest in the territory of north-western Europe (Cita, 1976; Brunnaker, 1979; Michaux et al., 1979; Lindsay et al., 1980; Rio et al., 1984; Kretzoi, 1987). Other authors (Azzaroli, 1977; Andreesku et al., 1987; Arrias et al., 1979, 1980; Pevzner and Vangengeim, 1986) infer an older, 3.9-4.1 Ma, age for the boundary. Both inferences are approximately equally supported by data, radiometric and paleomagnetic evidence inclusive, and both are debatable. The author considers all the three boundaries in mutual relations, as a particular reflection of one and the same global factor - climatic change. In fact, remains of mammals with Anancus, of probable Villafranchian age, occur above the reversed polarity beds, containing Globorotalia crassaformis in Marco Simone with a K/Ar age of 4.2 ? 0.2 Ma and a fission-track age of 3.32 k 0.3 Ma (Azzaroli, 1977; Arrias et al., 1979). Normal polarity beds, containing the Csarnotian fauna complex, in La Juliana, are underlain by a marine sequence with GI. crassaformis (Pevzner and Vangengeim, 1986). The Csarnotian complex, located in Wolferscheim (Boenigk et al., 1974), was also collected from the normal polarity beds, characterized by flora of the Brunssumian phytochron. In the two cases last mentioned the beds, characterized by the Csarnotian fauna, can be assigned to both SCT 17 (the base of Gauss) and SCT 21 (Nunivak event), since FAD of GI. crassaformis in the Atlantic is estimated at 4.3 Ma and their migration to the Mediterranean may have taken place during first tens of thousands of years. The Marco Simone beds may be also assigned to both SCT 16 and SCT 18. Repenning and Feifar (1976) estimate the time interval of Csarnotian fauna from 4.5 to 3.6 Ma. The author considers the three above boudaries in the marine and terrestrial sequences as facies manifestations of the common paleoclimatic boundary - a transition from the “almost non-glacial” conditions to a new long-term activation of glacial regime. The transition proceeded in three steps, recognized as kryo-SCT 20, 18 and 16. In bio-physiognomistic features the entire interval is transitional from Zanclean to Piacenzian, from Ruscilian to Villafranchian, and from Brunssumian to Reuverian. Therefore the author doubts that an unquestionable paleontological criterion for the solution of the discussed question about the Lower/Middle Pliocene boundary can be obtained, especially in regional plan. From the viewpoint o f global paleoclimatic periodicity and rhythmic cyclic recurrence with an interval of 3.7 Ma (see Chapter lo), it is advisable to draw this boundary near the SCT 21/20 transition. Floristic and faunistic findings from Sete, south of France, where SUC (1985) recorded the first signs of aridization in western Europe - the disappearance of taxodium species from the composition of forests - may be related exactly to SCT 20 (Michaux et al., 1979). Hence, it is wise to draw the boundary between the Lower (warm) and Middle (temperate) Pliocene in the Mediterranean area and in western Europe, like in the Pontian -Caspian basin, from the first surge of aridiza-

380

tion and cooling - Sete - Aidar - Karlaman, i.e. below kryo-SCT 20, at 4.05 - 4.01 Ma. This solution would be in agreement with data on the change in global atmospheric circulation, inferred by Stein (1986) from analysis of cores from holes drilled in the Northern and Southern Hemispheres. For more than a hundred years scientists have tried to tie in the question of the “first” appearance of “northern guests” - representatives of the North Atlantic marine fauna in the Mediterranean - to the problem of the Pliocene/Pleistocene boundary. The invalidity of such an approach was discussed in Chapter 3. Here we are interested in a purely paleoclimatic aspect of the problem, related to the questions whether the appearance of “northern guests” in the Mediterranean reflects global surges of cooling and what number of such appearances occurred in the Pliocene. Arctica islandica and Hyalinea baltica were traditionally regarded as the first “northern guests”. More than a hundred years ago from the presence of Arctica islandica Doderlein recognized the Sicilian, while the findings of Hyalinea baltica allowed Gignoux to recognize the Calabrian (Selli, 1977). Ruggieri and Sprovieri (1977) showed that the stratotypes of both stages are equal in extent and the base of both is dated roughly at 1 .O - 1.15 Ma. This implies that an attempt made at the IGC in London, 1948, to lower the Pleistocene boundary and place it below the base of Calabrian strata was erroneous. At that time the appearance of the cooler water ostracod Cyteropteron testudo was proposed to be regarded as the “first” appearance of “northern guests”; results of reliable joint investigations show that the ostracod Cyteropteron testudo appeared in a section at Vrica somewhat higher than sapropel “e”, whose age is dated at 1.64 Ma (Colalongo et al., 1982). Tenacious efforts of the Pliocene/Pleistocene Boundary Working Group made the IUGS Stratigraphic Commission recommend to accept this level as a new Pliocene/Pleistocene boundary (Aquirre and Pasini, 1985). However, this time again the recommendation to lower the Pleistocene boundary proved to be theoretically unsubstantiated. Ruggieri himself found that in an other section C. testudo appeared earlier, as compared with the section at Vrica. Previously, Arrias and co-workers (1980, 1984) and Bedini et al. (1981) proved that the boreal mollusc Arctica islandica had been present in Italian sections at least from the Reunion event, at 2.03 Ma. North Atlantic Globorotalia inflata also appeared in the Mediterranean about 2.2 Ma (Cita, 1976) to 2.05 Ma (Rio et al., 1982) BP. At last, studies performed by Spaak (1983) and Driever (1984) showed that the Mediterranean marine fauna yields a still earlier “northern guest” - foraminifer Fig. 11.11. Palynostratigraphic units of Pliocene Columbia and their correlation with key sections in Europe and in the ocean (after Hooghiemstra, 1984). A - North-western Europe July temperature changes (after Zagwijn and Doppert, 1978); B - Funza lake section (High plain of Bogota, Columbia, 4”50 N - 74”12 W , alt. 2547111). (a) pollen zones (1 55); (b) the depth scale (m); (c) age estimating - interpolation by fission track and K/Ar dating of ashes; (d) curve of the percentage of arboreal pollen. The intersections of this curve with the vertical dotted line indicate the moments in which the upper forest line passes through the high belt of the High plain and corresponds with a temperature of 10°C at this altitude (i.e. left peaks of the curve correspond to kryomers, right peaks correspond to thermomers). The dotted line is situated o n the 40% arboreal pollen level during the upper pollen zones 2 - 11, on the 35% arboreal pollen level during the pollen zones 15 - 30 (a correction is applied for the absence of Quercus (oak)) and o n the 20% level during the lower pollen zones 32 - 55 (a correction is applied for the absence of Quercus (oak) and Alnu.5 (alder)); C - oxygen-isotope record V28-239, the Pacific Ocean (after Shackleton and Opdyke, 1976). The numbers on the correlation levels, on the right, indicate isotope stages after Shackleton and Van Donk. -

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382

Neogloboquadrina atlantica, inhabiting the Mediterranean for a very narrow time span from 2.36 to 2.2-2.0 Ma. Thus, the appearance of “northern guests” in the Mediterranean marine fauna is recorded in SCT 6, 8, 10 and 12. Boreal Globorotalia crassaformis appeared in SCT 16 and, probably, SCT 18 and 20, while GI. conomiozea appeared in SCT 30. This is in agreement with climatic changes revealed by Thunell (1979) from processing by means of factor analysis of oxygen-isotopic and faunistic data (Table 11.2) from cores of Holes 125 and 132 drilled from the vessel Glomar Challenger. Let us now direct our attention to data on north-western Europe, generalized in column 6 (Fig. 11.8). They are widely known from studies by West (1968), Van der Hammen and co-workers (1971), Zagwijn (1974, 1985) and others. Column A (Fig. 11.1 1) illustrates the amplitude of summer temperature variations for 18 upper superclimathems. Characteristics of the Reuverian and Brunssunian are also presented in Fig. 11.8 after Meyer (1981). The brief review of the western European Pliocene shows that the climatostratigraphic principle of subdivision enables us to obtain results similar to Table 11.2. Water temperature estimation for the Mediterranean, obtained with the help of factor analysis of plankton foraminifera assemblages from DSDP site 132 core. Estimates are derived from the curves by Thunell (1979). ~~~

Zone

Temperature on the peaks of the curve (“C)

Correlation with superclimathems after the author

=,

T,”,,,,, 23 26 24 25.5

13 14 15 15

SCT SCT SCT SCT

Globorotalia inflata - Neogloboyuadrino pachyderma MPI 6

23 25.5 20 24.5

14 14 14 14

SCT 10

-

Stirone-Le Coupeet

SCT 11

-

Saint Vallier

Globigerinoides elongatus - Globorotalia crassaformis MPI 5

21

12

SCT 12

24 20 25 23

14.2 11.7 13.5 13

SCT SCT SCT SCT

Sphaeroidinellopsis subdehistens

26

14.5

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?

Gfoboroialia punctirlata MPI 3

? 23.5 26

? 13.5 15

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?

Globoroialia margariiae MPI 2

21 24.5 22 24

12 14 13.5 14

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23

13.5

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lntrl

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Acquitraversan Montopoli

-

Etouaires? Triversa

?

383

those for the Pontian-Caspian, i.e. we can recognize the same number of units, whose age and position relative to the paleomagnetic scale prove to be similar in all three regions. Now we continue brief correlation with some other regions. Column B (Fig. 11.1 1) presents results of the unique investigation carried out by Hooghiemstra (1984) on an upland in central Colombia. The reader is well acquainted with the 1 1

2

3

4

5

6

8

7

I

u 1 . 1

29

Fig. 11.12. Time units of the continental Pliocene of ihe Northern Hemisphere according to mammal fauna. 1 - paleomagnetic scale (age (Ma)) (after Berggren et al., 1985, McDougall et a!., 1984) with corrections(*); 2 - Europe (after Mein, 1975); 3 - Africa (after Berggren et al., 1985h; Khrisanfova, 1987; Harris and Johanson, 1986); 4 - India (after Lindsay et al., 1983); 5 - Northern Kazakhstan and the southern part of West Siberia - author's summary based on the data of Zazhigin (1980). Zazhigin and Zykin (1984), Zinova (1982). Zykin (1986), Zykin and Zazhigin (1987), Marrynov et al. (1987). Shkatova (1987); 6 - Lake Baikal region, Mongolia arid north-eastern Asia - summary based on the data of Bazarov (1986), Devyatkin (1981). Zhegallo (1978, 1985). Pevzner et al. (1982), Shilova (l981), Sher and Kaplina (1979); 7 - North America (after Tedford, 1981; Berggren et al., 198Sh; Repenning and Fejfar, 1976; Repenning, 1987; Macfadden, 1979; Neville et al., 1979); 8 - superclimatherny and hyperclimathmes, their calculated age.

384

pollen record of Pleistocene lacustrine sequences, studied by Van der Hammen. Now his pupil analyzed an upper, 357 m long, core from an 800 m section of a former lake in the vicinity of Bogota, at a height of 2547 m. The plateau is surrounded by mountains with forested slopes, but the upland is treeless. Pollen rain well reflects vertical shifts of vegetation zones and, ultimately, changes in temperature and atmospheric precipitation. During warming and humidification the forest zone widens, during cooling and aridization the zone becomes narrower, but the upper herb layer widens. Sedimentation in the lake was continuous throughout the Pliocene. The presence of numerous ash layers allows dating of the section by the K/Ar and fission-track methods (1 3 datings). All in all, 26 glacial - interglacial cycles, consisting, as a rule, of two phases (kryomer and thermomer), were recognized in the interval of 3.6 Ma. The absence of paleomagnetic datum in the Funza section does not allow us to consider the age estimates as absolutely reliable. Therefore its correlation with European schemes differs in details. Nevertheless data on the Funza section brilliantly confirm the existence of global climatic cycles related to a 400 ka rhythm of eccentricity and practical importance of superclimathems as the main climatostratigraphic units for the Pliocene. The climatostratigraphy of the Pliocene of other regions is poorly developed, although some strides have been recently made towards its deciphering, as seen from Fig. 11.12 and 1 1.13. Fig. 11.12 collates biostratigraphic schemes for six regions, namely, East Africa (column 3), the Irtysh River basin (column 5), central and north-eastern Asia (column 6 ) and North America (column 7). All of them are based on the evolution of mammals (and freshwater molluscs in the Irtysh River basin) and are fairly well comparable. Six stages of fauna evolution, climatic in nature, can be recognized for at least two of these. Fig. 11.13 generalizes events of the Pliocene history of the kryosphere and the ocean. There are recognized three main stages of: (1) asymmetrical development of the kryosphere, when glaciation in the Southern Hemisphere attains its maximum; (2) its substantial degradation; (3) pulsating continental glaciation of the Northern Hemisphere. Judging from radiometric data, multiple activations of glaciation are related to the chronological succession of superclimathems. The second stage is associated with the great Middle Pliocene ocean transgression (Azores - Kanary), which reached 25 - 30 m in height at about 4.2 Ma. Its glacio - eustatic character is almost undoubted. 11.4. The main steps in the Pliocene climate evolution

In the time interval from 7.15 to 1.0 Ma, i.e. from the base of the Pontian to that of the Chaudian, 30 superclimathems have been identified, which embrace 10 hyperclimathems (Fig. 11.14). These superclimathems have been distinguished as a result of interpreting the data on the deep-sea drilling as well as from 18 regional sequences on the continents (the Black Sea, the Ukraine, the Caspian Sea, the Ural foothills, Byelorussia, the Altai Mountains, the north-east of the USSR, Mongolia, the Mediterranean, the north-west of Europe, England, Iceland, the Pacific coast of North America, the Middle West of the USA and the Atlantic coast, New Zealand, Patagonia, Antarctica and Java). The basic data considered were those from the Black Sea, the Ukraine and the Caspian Sea known to the author from

385

Fig. 11.13. The evolution of cryosphere and ocean levels change in the Pliocene. I - Paleomagnetic scale (after McDougall et al., 1984 and Berggren et al., 1985) with corrections; 2 - glaciation in Antarctica (after Mayewski, 1975; Ludwig et al., 1980; Leonard et al., 1983; Ciesielski et al., 1982) and Patagonian Andes after Mercer, 1976; Mercer and Sutter, 1982); 3 - New Zealand marine events o n the North American coasts: ( a ) along (after Loutit and Kennett, 1979, 1981); 4 Santa-Rosa Islands (Ridge Southern California) (on the left) (after Crouch and Poag, 1979), (b) along the Atlantic coast (after Blackwelder, 1981 and Cronin et al., 1985), (c) the Azores and Canary Islands transgression (after Meco and Steams, 1981); (d) mammalian fauna exchange between North and South America (Uquian and Chapadmalal ages) (after Berggren et al., 1985b); 5 Iceland (after Einarsson et al., 1967; Gladenkov, 1978; Zagwijn, 1974) and Faeroes-Greenland Ridge (after Shor and Poore, the western part of the Arctic basin (after Zarkhidze, 1981; Baranovskaya and Zarkhidze, 1981); 6 1987; Funder et a]., 1985) and Alaska coast (after Hopkins et al., 1973; Denton and Armstrong, 1969; Carter et al., 1986); 7 - superclimathems and hyperclimathems, their calculated age. ~

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386

his personal field investigations and the published data on the Mediterranean, north-west Europe and North America. This information is summarized in Figs. 11.1, 11.5, 11.8, 11.12 and 11.13. The marine section of the Black Sea Late Miocene and Pliocene (Kerch - Tamanian region and Georgia) and soil - loess section of the southern Ukraine (the Crimea and the Azov basin) are adjacent to each other, which makes it possible to observe the lateral facial alternation of marine and subaerial formations and to make a summary section of the Black Sea with the help of palaeomagnetic data. The sequence of the Black Sea palaeoclimatic events, which is unique in its completeness and includes four stages (Fig. l l . l ) , and a relatively large number of corresponding palaeontological and palaeomagnetic studies allow us to consider it as a possible standard for developing a general climatostratigraphic scale of the last 7 Ma interval of geological history. 11.4.1. Palaeoclimates of the Early Pliocene, 7.15 - 4.7 Ma

For the 7.15 - 4.7 Ma time interval, only hyperclimathems can be globally observed. Superclimathems are revealed in the well-studied regions - in the Black Sea area SUPER- AND

HYPERCLIMATHEMS OF THE LAST

A Magnet hems

-R

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0

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summer

:-%,

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Likhvin

-

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Mammalian ages

8

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7 M. YR.

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5

Ti r o s p o l ion

71.0

Tam a n ia n

_--__--

Psekups Khapry NM 17

= _ _ - Kvabebi

-

NM 16

> / Ruscinian

NM 15

-

-

NM 14

NM 13

Turolian

- NM 12

-Fig. 11.14. Tentative generalizing curve showing the changes of the suggested mean temperatures of the warmest month on the areas between 4 5 " and 60" N (for the last 7 Ma) (after Zubakov and Borzenkova 1988, with corrections.

387

(Fig. 11.1) and in western Europe (Fig. 11.8) as well as in New Zealand (Fig. 11.9). This long time interval can be distinguished as the time of the maximum glaciation in the East Antarctic and the onset of the West Antarctic sheet glaciations (Fig. 1 1.13). The latter is dated as 7.4 Ma by K/Ar according to a palagonitic formation in the Hadson and Jones Mountains and the maximum Queen Maud Land glaciation which is estimated from 5.5 to 4.5 Ma (Mayewski, 1975). According to independent geological records and isotopic data the volume of the Antarctic glaciers at that time was 1.5 - 2 times greater than at present (Shackleton and Kennett, 1973). That was the time of a sharp climatic asymmetry of the Southern and Northern Hemispheres (Flohn, 1974). The most detailed climatostratigraphic division of this long interval is revealed in the Black sea basin in the Capes Panagiya - Zhelezny Rog section on the Tamanian Peninsula (Fig. 1 1.2) by alternating freshwater beds (kryomers) and brackish layers (thermomers). The Messinian salinity crisis, 6.4- 5.3 Ma, corresponds to the XI and X HCTs with climate alternating from a cool to a warm dry one in the Mediterranean area. The cause of the drastic evaporitic sedimentation is open to discussion. Some scientists associate the cyclic recurrence of evaporites with periodic isolation of the Mediterranean from the oceans brought about by plate motions which resulted in evaporite drawndown and eventual desiccation (Ryan, Hsu et al., 1973; Ryan et al., 1974; Cita, 1979; Cita et al., 1973). Others suppose the existence of difficult but continuous connection between the Mediterranean and the Atlantic (Stanley and Wezel, 1985). It is possible that the ten evaporitic cyclothems of the Messinian in the sections of the North Apennines indicated by G.B. Vai and F.R. Lucchia (Catalano et al., 1978, pp. 217 -249) represent climatic events of the same rank that the orthoclimathems of the Pleistocene with the duration of 80,000 to 120,000 yr belong to. In the Black Sea the Messinian is represented by a series consisting of 5 or 7 climathems (Fig. 11. l), which are dated by the palaeomagnetic method in the sections of Capes Panagiya-Zhelezny Rog, the Tamanian Peninsula (Fig. 11.2). The 35th, 34th and 33d SCTs correspond to the Early Pontian (Novorossian substage) and are correlated with the Tripoli and Terravechia formations in the sections in Sicily (Catalano et al., 1978; Cita, 1979), i.e. with the pre-evaporitic Messinian. The Middle Pontian (Portafarian), i.e. SCT 32, marked by a cooling trend and an invasion of Pannonian waters with Congeria subrornboides, coincides with N-zones within the 6th epoch of polarity (6.4 Ma). This is the most important climatic boundary recorded globally. It is associated with a dramatic shift in 6 I3C by O.8%0 in the Pacific (Loutit and Kennett, 1979; Keigwin and Shackleton, 1980), with the uplift of the isthmus of Gibraltar and the onset of the Messinian salinity crisis as well as the culmination of the East Antarctic glaciation (Ryan et al., 1974). The upper Pontian (Bosphorian) is divided into two horizons. The lower horizon corresponds t o the Znamenka thermochron (Veklich, 1982) and the lower evaporites o f the Mediterranean (Cessoso-solifera formation). According to the data of Catalano et al. (1978), the latter has three subunits and therefore it corresponds to the 31st, 30th and 29th SCTs. The upper horizon, the Belbekian (Veklich, 1982), has a reliable correlation with the upper evaporites and the Lago Mare formation of the Mediterranean (Fig. 11,8). The fact that they are of the same age is confirmed by the presence of a common Late Bosporian association of brackish molluscs with

388

Congeria subcarinata and Limnocardium aff. inlongeavum and ostracods Cyprideis pannonica, C. agrigentina and other species. According to Sissingh (1972), and Cita and Colombo (1979) this complex spread from the Black Sea to the Mediterranean right up to the shores of Spain. The Belbek cooling of the 28th SCT is confirmed by the oceanic regression which can be dated as 5.5-5.6 Ma in the Bou Reg Reg section in Morocco (Cita and Colombo, 1979). The standard of the 27th SCT in the Black Sea appears to be the lower ore layers of the Azovian horizon of the Kimmerian stage with Paradacna deformis. The termination of the Messinian Age is marked by an almost complete desiccation of the Mediterranean and the Black Sea. This event is manifested (after Hsii and Giovanoli, 1980) by a horizon of algae stromatolites and gravels at a depth of 864 m in the drilling section DSDP-380 in the Black Sea. In the Mediterranean it corresponds to the sedimentation of Arenazzolo sands and the formation of underwater canyons (Cita and Colombo, 1979, Chumakov, 1974, 1982). 11.4.2. The Middle Pliocene warm climates, 4.7-3.65 Ma The standard type of the Middle Pliocene is the Brunssumian regional stage, which is clearly marked in a new section at Wittmunder Forest, West Germany (Meyer, 1981), on the North Sea coast and splits there into five or more superclimathems (Fig. 11.8). In the Black Sea basin the Middle Pliocene is represented by the Kamyshburunian Horizon reflecting the increasing salinity. It contains three beds of oolitic iron ore, the lower one being included into the HCT VIII. The subaerial standard of the HCT VIII is the Sevastopolian pedocomplex comprising 5 to 10 buried soils of the savanna type (Veklich, 1982; Sirenko and Turlo, 1986). Their analogue in North America (Fig. 11.12) is probably the upper layers of the Early Blancan beds (Repenning and Feifar, 1976). In the Mediterranean the Middle Pliocene warming is represented by the Tabianian stage with Ficus ficoides (Br.) and Globorotalia margaritae evoluta (Berggren and Van Couvering, 1974), which corresponds t o the middle of zone N 19 and the zone MPI, (Cita, 1976). The hyperclimathem VIII, in particular SCT 21, is synchronous with the greatest Pliocene rise of sea level reaching, according to different estimates, from 28 - 36 m to 60 m (Blackwelder, 1981). Its age determined by K/Ar measurements on the Azores and Canary Islands is 4.4 to 3.7 Ma (Meco and Stearns, 1981), on the Atlantic coast of the USA, 4.5 & 0.2 Ma and in the Aegean Sea, according to Benda, 4.4 to 3.6 Ma (Blackwelder, 1981). This transgression (Fig. 11.13) caused the opening of the Bering Strait and the distribution of the Pacific molluscs Serripes groenlandicus and other species across the Polar basin to the Norwegian and North Seas (Herman and Hopkins, 1980). In the Iceland sections, the appearance of Pacific molluscs is referred to the sub-zone Yoldia myalis (Gladenkov, 1978) which is conventionally comparable with SCTs 17 - 19 (Fig. 11.13). The Mactra zone corresponds to SCTs 21 - 23 (or 21 - 27) as well as the Pecten and Scallop Hill beds on the shores of the Ross Sea and Black Island, Antarctica, whose age is estimated by K/Ar dates as over 4.2 and 3.85 Ma (Mayewski, 1975; Mercer, 1978). The climate of the Middle Pliocene, particularly of the HCT VIII, was on the whole the warmest

389

of the last 7 Ma. However, it was not homogeneous. In the Southern Hemisphere the conditions of the 21st SCT, the Pecten time-Solo IV time (Figs. 11.15, 11.16), were warmest. During this time (the Nunivak event and a bit earlier) the summer temperature of the surface water near the shores of the Antarctic continent at latitudes from 57" to 69" was 7" to 10°C higher than at present (Ciesielaki and Weaver, 1974). The calculation was made taking account of variations in the ratio of the warm-water silicoflagellates genus Dictyocha to the coldwater Distephanus. During the 20th SCT the temperature was decreasing to the present level and during the 19th SCT, which coincides with N-event Cochiti, it was 3" to 7°C higher than the present temperature. During the 18th SCT, corresponding to the Lago Viedma glaciation in Patagonia, 3.55 Ma by K/Ar according to Mercer (1978), the warm water Dictyocha was not found at all in the samples of "Eltanin". A similar picture is revealed in the Northern Hemisphere, where the 21st SCT was also warmest with the early Csarnotan and Cherlak fauna (Fig. 11.12). SCTs 19- 17 with the Late Csarnotan fauna were Iess warm. According to palynological data (Meyer, 1981), the 19th SCT was even colder. There are no doubts that the West Antarctic ice sheet completely degraded in SCTs 23, 21, 19 and 17. Moreover, the presence of the Pecten beds on the shores of the Ross Sea, the isotopic data (Fig. 11.15), bioclimatic records (Fig. 11.16, 11.17) and high sea level (up to 28- 36 m) in the tectonically stable parts of the

1

2 4.5 4.0 3.5'4. Cold

Warm

3

t' winter C

5

77.5 725 t3.5

29 N: pachydcrma factor (cold)

Fig. 1 1 . 1 5 . Possible correlation of oxygen-isotope and bioclimatic curves (composed as a result of studying faunistic data by the factor analysis method) based on comparing curve peaks with superclimathems sequence. 1 - Oxygen-isotope curve of benthic foraminifers Cloboeassidulina subglobossa (core V28 - 179, equatorial zone of the Pacific Ocean) (after Shackleton and Opdyke, 1977); 2 - oxygen-isotope curve Mediterranean sur(DSDP 397 site, equatorial Atlantic Ocean) (after Shackleton and Cita, 1979); 3 face water temperatures deduced from planktonic foraminifers complex (DSDP 132 site) (after Thunnell, 1979); 4 - Polar front boundary shifts in the South Atlantic deduced from planktonic foraminifers complex (DSDP 514 site) (after Ludwig et al., 1980); 5 - relative fluctuation of water temperature in the Southern Hemisphere high latitudes determined by the factor analysis method (river Blind section, New relative water temperature changes near Java coasts (the Zealand) (after Loutit and Kennett, 1979); 6 Solo River section) (after Van Goersel and Troelstra, 1981). ~

~

MAGNETIC POLARITY

EPOCHS

cLwmmAn-K: ZONES SOLO R I V E R SECTION

p:,","[z",:c

FORAM ZONES

MEDITERRANEAN STAGES

Olduvai

2

COLU

MATUYAMA

SANTEHNIAN I Basal CALA6RIANs.l.)

Global Palaeoclimate of the Late Cenozoic (Developments in Palaeontology and Stratigraphy)

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Global Palaeoclimate of the Late Cenozoic (Developments in Palaeontology and Stratigraphy)

Late Neogene (Developments in palaeontology and stratigraphy)

Global Biogeography (Developments in Palaeontology and Stratigraphy)

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Quaternary Dating Methods (Developments in Palaeontology and Stratigraphy 7)

Terminal Eocene Events (Developments in Palaeontology and Stratigraphy)

Proxies in Late Cenozoic Paleoceanography (Developments in Marine Geology)

Fossil Nonmarine Ostracoda of the United States (Developments in Palaeontology and Stratigraphy)

Patterns of evolution, as illustrated by the fossil record, Volume 5 (Developments in Palaeontology and Stratigraphy)

Evolution and Geological Significance of Larger Benthic Foraminifera (Developments in Palaeontology and Stratigraphy, V21)

Permian-Triassic Evolution of Tethys and Western Circum-Pacific (Developments in Palaeontology and Stratigraphy)

Permian-Triassic Evolution of Tethys and Western Circum-Pacific (Developments in Palaeontology and Stratigraphy)

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Stratigraphic Micropalaeontology of Atlantic Basin and Borderlands (Developments in Palaeontology and Stratigraphy)

Mesozoic and Cenozoic sequence stratigraphy of European basins

The Late Cenozoic of Patagonia and Tierra del Fuego (Developments in Quaternary Sciences, Volume 11)

The Jurassic and Cretaceous stratigraphy and palaeontology of Traill Ø, East Greenland

Palaeontology of California. Philadelphia

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Cenozoic Mammals of Africa

Principles of Sedimentology and Stratigraphy

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Principles of Stratigraphy

Principles of Sequence Stratigraphy

Insider Trading: Global Developments and Analysis

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Lunar Stratigraphy and Sedimentology (Developments in solar system- and space science)

Principles of Stratigraphy

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