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 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

xi

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

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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

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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

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PART I

THE GLOBAL CLIMATIC EVENTS OF THE PLEISTOCENE

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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).

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SECTION I

METHODOLOGICAL PROBLEMS OF PALAEOCLIMATOLOGY

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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.

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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

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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

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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).

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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

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I

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40

1

1

20

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0

I

I

20

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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.

This Page Intentionally Left Blank

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 .- - -

<

- -3417-

beudoemLLian Lacunosa

_ _ _ _ - -4 i Fig. 3 . 1 , Stratigraphic framework of the Eastern Mediterranean deep-sea record. The first two columns on the left show the isotopic curve obtained after Globigerinoides ruber in core R C 9 - 181 and the columnar log of the core, which contains 12 discrete sapropels. Column 3 shows the distinction of climatic zones (of Ericson and Wollin, 1968) and column 4 the nannofossil biozones (of Gartner, 1977). From Parisi and Cita (1982).

73

Grotte Le Vallonnet, near the town of Mentona, French Riviera, where effects of rather warm and humid climate were found by Renault-Miskovsky and Girard (1978) in members E and D. In many works it is named the Calabrian. The theromeric phase of subcycle F, (Giinz - Mindel) includes Portuensian transgression, established on the Lazio coast (Ambrosetti et al., 1972) and, probably, levels U,, 75 - 83 m , and U,, 60- 65 m, at Mallorca. The Vallonnet climatozone (or its lower part) comprising representatives of macchia ( = gariques) with porcupine Hystrix) may be also assigned to the same phase (Chaline, 1977, 1978). According to Liimley, sediments synchronous to the cycle at the Grotte Le Vallonnet have normal polarity and are associated with the Jaramillo event at about 0.97 - 0.9 Ma (Renault-Miskovsky, 1978). The kryomeric phase consists of the Ficarazzi Clay in the type section of the Sicilian at Cava Puleo of Ficarazzi, where it is characterized by the appearance of cold water pteropods Limacina retroversa and the like (Ruggieri et al., 1984). In the Grotte Le Vallonnet section the phase includes members B, and B,, whose normal polarity may be related t o an inferred normal polarity event at 830 ka or to the Jaramillo event (Renault-Miskovsky and Girard, 1978). This kryophase is believed to be associated with Gunz gravels in the Ero Valley, France, which underlie pyroclastics of the Agde volcano with K/Ar age of 740 ka (Monjuvent et al., 1984). Subcycle F, is correlative with isotope stages 25 -22 (Ruggieri et al., 1984). According to Azzaroli (1983), cycle F starts the Galerian phase in the evolution of mammals. Subcycle F, includes the lower level of raised beach complex U (after Butzer) 30- 33 m Flaminian ( = Mindel I) kryochron that could be correlative with G r k e and Escale bioclimatozone A - F (Chaline, 1977). This cooling consisted of several phases which witnessed the development of steppes inhabited by arcto-steppe rodent assemblages on the northern shore of the Mediterranean Sea. The age of the cooling determined on pyroclastics from deltaic gravels of the Portuensian regression falls in the range of 680-706 & 70 ka (Ambrosetti et al., 1972, 1978). Cycle E started with the Milazzian' transgression corresponding to raised beaches V , and V2 on Mallorca. In the continental section the standard of this thermomer can be represented by Escale ( = St. Esteve) bioclimatozone G whose fauna contains monkeys Macacus florentinae and water voles Arvicola cantiana, while pollen and spore assemblages suggest the distribution of forests with abundant Celtis (Bonifey, 1980). The cycle ended with the Nomentanan kryophase (Mindel-II?) with K/Ar ages on tuff with black pumice from Bracciano Volcano at Torre in Pietro and other sections in the range of 438-417 ka (Ambrosetti et al., 1978). In the continental section the kryophase is characterized by the final Galerian fauna Ranuccio with a K/Ar age range of > 365 k 4 to < 487 f 7 ka (Azzaroli, 1983). Pollen complexes o f peat bogs of that time point to a drastic depletion in floristic composition of forests and the extinction of thermophilic species (Ambrosetti et al., 1972). The appearance of lemmings in bioclimatozone Escal H - (Chaline, 1977) and carnivores, such as Vuipes praeglacialis and Canis etruscus rnosbachensis (Bonifey, 1983), as

' O r Tarquinian after Arnbrosetti et al. (1972,

1978)

well as Ursus arctos, Cuon cf. alpinus and Capreolus capreolus in the Ranuccio (Azzaroli, 1983) suggests the development of arcto-steppe landscapes directly on the Mediterranean shore of France. It should be noted that this cooling put an end to the Galerian stage in the evolution of mammals which lasted from 1.O - 0.9 approximately to 420 ka BP. The cooling also put an end to the evolution of marine nannoplankton, since the disappearance of Pseudoemiliania lacunosa in deep-sea Mediterranean cores is dated at 430 ka BP (Raffi and Rio, 1979; Parisi and Cita, 1983). Cycle E can be correlated with isotope stages 15 - 14 or 15 - 12. Warming related to cycle D is marked by a marine transgression which left levels Wl,2,3 at Mallorca (Bowen, 1978), beds with Patella ferruginea, 23 m level, at Grotte Lazaret, underiain by an Early Acheulian cuItural layer (Liimley, 1968). The Tarquinio raised beach of the Tyrrhenian Sea, whose stratigraphic position is uncertain, could also be assigned to this warming. Its continental equivalent is represented by the Pariolian beds with forest fauna including Palaeoloxodon antiquus and Hippopotamus amphibius, as well as Via Flaminia tuff with macrofossils and pollen suggesting the distribution of forests with rich specific composition, including Abies, Bums, Celtis australis, Fagus silvatica, Pterocarya massalongi, Zelcova crenata, Taxus baccata, many species of Laurus, Quercus, and the like (Ambrosetti et al., 1972). In Rodentia the warming is marked by the migration of porcupine Hystrix - a typical inhabitant of subtropical macchia - towards northern France (Chaline, 1977). Subcycle W - “paleo-Tyrrhenian”, if it could be called so (after Issel in Zeuner, 1959), is correlative with isotope stages 13 and 11. Subsequent warming related to the Riano sequence and zone Orgnac 3G is poorly recorded. The spore and pollen diagram of the zone below the tuff level of Riano (280 f 30 ka K/Ar age) gives evidence of an oceanic temperate forest with dominant thermophilic species, such as Pterocarya, Guercetum mixtum, Zelvoca, etc., followed by an oceanic cool forest above the tuff level (Ambrosetti et al., 1972). Cycle D ended with the Ostian erosional phase and La Fage bioclimatozone correlative with isotope stage 8. Cycle C is marked by the first appearance in the Mediterranean Sea of tropical Senegalese molluscan fauna dominated by Strombus bubonius and coccolithophorid Emiliania huxleyi as well. The Strombus I zone recorded in raised beach XI, Mallorca, is dated by U-series at 210 k 10 ka (Bowen, 1978). Erniliania huxleyi occurs directly above the top of the lower sapropel, which is one of the thickest, in cores through the eastern Mediterranean dated at 225 - 220 ka (Parisi and Cita, 1982). This Eutyrrhenian or, after Zeuner (1 959), Monastirian, transgression marked in Provence by a 15 - 18 m raised beach (Liimley, 1968), left signs at two levels. The second and higher level of transgression X, dated by U-series (Mallorca) and by fission-tracks (Italy) yields are ranges of 190 f 10 ka (Bowen, 1978) and 177 f 30 ka (Ambrosetti et al., 1972) respectively. The level is characterized by poor Senegalese fauna; it could be just this level which was described by Liimley at Grotte Prince as a 15 m high “inter-Riss” raised beach. Bipartition of the Eotyrrhenian transgression and radiometric datings unambiguously indicate a valid correlation with isotope stage 7. On land the Eotyrrhenian thermochron is equivalent t o the Grimaldi bioclimatozone with Late Acheulian cultural layers. Cycle C ended with the Fontechevade climatozone and is correlative with isotope stages 7 and 6.

75

Cycle B - Neotyrrhenian or (after Zeuner, 1959) Late Monastirian - has been studied at greater length, as compared with all those discussed above. It comprises three raised beaches. The highest (10- 12 m) raised beach (Grotte Raphael, level NORTH A T L A N T I C

M A C E D O N I A

POSTGLAC IA L ~

I

1

6

-

C

m

7b:

PA N G A I 0 N

I

I-+ I

B

SYMVOLON

]

I

D

lx

9 10

t

E T

I1

LEKANIS COMPLEX

J

12

F 13 14

m G

15

1!'

B O Z DAGH COMPLEX

m

16

H 17

PHALAKRON

m

Fig. 3.2. Mediterranean - North Atlantic correlation (after Kukla, 1977, fig. 3). Foraminifera1 assemblages in North Atlantic core K 708 - 7 after Ruddiman and McIntyre (19761, compared with the pollen record in the Tenaghi - Phillipon peat bog in Macedonia, after Van der Hammen et al. (1972). The percentage of subtropical and transitional foraminifera in K 708 - 7 and oak pollen in Macedonia are shown in black. Forams other than a polar type and total tree pollen stippled. Position of terminations, correlation with oxygen isotope stages and glacial cycles modified after Ruddiman and McIntyre (1976).

76

*

Y1) is characterized by a typical Senegalese fauna and dated by U-series at 125 10 ka (Bowen, 1978), and by fission-tracks at around 127 k 13 ka (Ambrosetti et al., 1972). In time it corresponds to the formation of the thickest sapropel S5 with an age of 125 - 116 ka estimated for 15 cores recovered by R/V Lynch and Chaine in the Strait of Sicily (Muerdter and Kennett, 1984). The age of the second level, 2 - 3 m high (Grotte Madona del Arma, level Y2), known as “partly Strombus” ranges from 110 -t 5 to 95 k 5 ka as dated by the U-series method, and is about 90 18 ka as determined by the fission-track technique. The age is in agreement with that for thin sapropel S4 at 99- 101 ka. The third, Y, level, 0.5-2 m high, referred to as Epityrrhenian, is dated at 80 f 5 ka on Mallorca. The level is synchronous with sapropel S3 dated at 82 - 79 ka (Muerdter and Kennett, 1984). The above data unambiguously indicate that the three Neotyrrhenian warm peaks COTrespond to isotope substages 5e, 5c, and 5a with ages of 126- 115, 105 -95, and 85 - 80 ka, respectively (Duplessy, 1980, 1981). Over the section of unique Tenagi Philippon peat bog, Macedonia (Fig. 3.2), which was thoroughly investigated with the aid of pollen analysis, substage 5e and sapropel S5 can be assigned to the Panagaion interglacial, when the mountains of northern Greece were covered by deciduous broad-leaved forests with dominant Carpinus and Quercus cerris. Substage 5c and sapropel S4 are associated with the Drama interstadial characterized by the distribution of deciduous forests with Fagus and Carpinus, while substage 5a and sapropel S3 are related to the Elevteroupolis interstadial (Wijmstra, 1978). Cycle B ended with the Pontinian kryochron and Loisia bioclimatozone, when numerous short-term climatic oscillations could be established.

*

3.2. The Caspian basin as a major record of changes in humidification in interior Eurasia A stratigraphic succession of sedimentary strata of the Caspian Sea (or Lake) is a natural stratotype for the subdivision of the Pleistocene of the entire Caspian drainage basin including the Volga, Ural, Kura river basins, and, sometimes, the Amu Darya and Syr Darya basins. Local stratigraphic schemes for the Pleistocene (and Pliocene) of the east European part of the USSR, the Caucasus, West Siberia, Kazakhstan, and Central Asia have been always tied in to the Caspian marine sequence. It is due t o this fact that the formal stratigraphic schemes of the Pleistocene and Pliocene of the USSR are based on the Caspian marker levels. The geologists studying the area have taken their cue from the history of the closed inland basin; this was one of the causes leading to an independent development in Russia of stratigraphic classification for Quaternary deposits; in particular, it concerns the Pleistocene -Pliocene boundary. In the USSR the Pleistocene - Pliocene boundary marked by the appearance of brackish water molluscs of currently living specific composition is drawn at the top of the Apsheronian and the base of the Baku regional stages of the Caspian basin. The Plio - Pleistocene history of the Caspian is characterized by numerous changes in sea level and associated variations in salinity from 5.7%0to 15 - 18%0 (at present, salinity is 5 -7%o in the Volga delta and 12- 13%0in the southern basin).

77

The present Caspian sea level is 28 m below sea level. During high transgressions absolute sea level reached 50 m, and the water area and volume roughly doubled (Svitoch et al., 1981). In the Manych Strait, whose elevation is 26 m, this resulted in unidirectional run-off of the Caspian waters into the Azov and Black Sea basin. The deepest buried lines have been recorded on the Caspian Sea floor at absolute depths of -40 to - 60- 80 m (Kalinin et al., 1976). Thus, as a result of deep regressions, the Caspian northern basin was completely dried out, and the total amplitude of changes in sea level reached 120 m. Until very recently, the genetic interpretation of sea level changes in the Caspian Sea has been a subject of controversy, for it was difficult to determine the contributions of climatic and tectonic components. Along the Caspian coast tectonic manifestations are impressive. South of Makhachkala, along the Caucasus coast, near-shore marine facies form about 15 abrasion and accumulation terraces 200 to 400 m high. The northern Caspian area and the Kura River lowland are occupied by a young Khvalyn marine plain having no more than three or four levels; at the base of the plain a Pleistocene sedimentary sequence with a normal stratigraphic succession, up to 1 km thick, is exposed. Datings of the marine terraces suggest that the rate of tectonic uplift of the Caucasus coast off the Caspian Sea could be 0.3 - 0.4 mm/y. The rate of eustatic changes in the Caspian sea level could be an order of magnitude larger. Using I4C and TL measurements, Svitoch and coChange of the Leuel, rn (Recent -28m abs.)

b

u

ATvp "C (Recent IO-IZ"C)

c

t)

APmm 400 200

Mount

qLacLat Lon -200

mv

..

,,,*:::::]

'-

n....

EarLy KhuaLyn

c

Late Khazar

c

x..-.

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

'/A

/

...?

...........

..'*...... ? x..'

c

.-,

'\

x. .. x'. ....

Duzdag

A-

\

-- -/

A/--

' A ' A

---

c

Fig. 3 . 3 . Caspian sea level change (after the author) and mean annual temperature and nieaii annual precipitation reconstruction for the western Caspian coast (after Abramova, 1982) and along the Paleolithic Azykh Site section (after Velichko et al., 1980). The Alpine glaciation of the Cauca\us shoun after Milankovsky and Koronovsky (1977). 1 - reversed polarity; 2 - normal polarity; 3 - run-off Caspian water through the Manych Strait.

78

workers (1981) and the author determined that the rate reached 5 mm/y during short-term intervals of several thousand years. The net rate of recession of sea level from the terrace of the maximum phase of the Khvalyn transgression, 48-52 m high (abs.), whose surface has not been deformed tectonically, to the post-Khvalyn regression, 40-42 m below sea level (Fig. 3.3) is 1 mm/yr. Although the climatic nature of the changes in sea level and salinity in the Caspian Sea seems evident to the majority of investigators, the mechanism of this relationship remains uncertain. On general theoretical grounds, some authors attribute the Caspian transgression to the melting of ice sheets covering the Russian platform (Fedorov, 1957; Moskvitin, 1962, Kvasov, 1975) and West Siberia (Volkov et al., 1978; Grosswald, 1983) and date them by the end of the glacial and the beginning of the interglacial phases. However, there is no agreement between the extent of continental glaciations and that of the Caspian transgressions. For example, the maximal (Khvalyn) transgression is synchronous with the minor Wiirm glacial, and the minor Khazar transgression corresponds to the maximal Dnieper Glacial. Empirical data indicate that the rises in the Caspian sea level are synchronous with kryomeric phases, or, to be exact, to their onset, while the falls in sea level are related to thermomeric phases, or their onset. This is evidenced from (i) facies changes: marine terraces located along the Caucasus coast gave way to fluvioglacial stream terraces of the Caucasus Range (Kozhevnikov et al., 1977; Vasiliev, 1984); (ii) overlapping of beds containing Caspian and Mediterranean marine molluscs in the Manych Strait (Popov, 1983); and (iii) pollen analysis. Thus Abramova (1982), who studied about 2000 specimens collected from terraces of the Caucasus coast, found out that steppe and semidesert pollen spectra are associated with the regressive facies, while spectra rich in pollen and spores of the plants typical of the forest zone are related t o the transgressive facies. Using pollen statistics, she estimated variations in annual temperature and in annual atmospheric precipitation for 12 Pleistocene time intervals (Fig. 3.3). Oxygen-isotope analysis of Caspian molluscan tests performed by Gorbarenko and Nikolaev (Svitoch et al., 1981) has shown that at the regressive phases 6 l 8 0 values varied from 1 - 0.4%0off the Caucasus coast to - 2.2% at the Volga delta. In the transgressive phases 6 l 8 0 values fell to - 4 - 6%0and even - 10 - 1 l%o;this can be explained only by a large river and glacial discharge into the Caspian Sea. Paleontological data for the southern Caspian basin (Fedorov, 1957; Alizade et al., 1972, 1978, 1984) are in a good agreement with the isotopic record. This suggests that the salinity of Caspian waters decreased to 5 - 1 l%oat the transgressive phases, and increased to 15 - 20%0at the regressive phases in the southern basin. According to Svitoch and co-workers (1981), in the northern basin the variations in salinity conditions must have been more complicated. Although the regressions caused the shallowness of the basin to 6 - 10 m, local low salinity conditions might have ocurred there due to a higher river run-off (Rychagov, 1977). The above rapid variations in salinity conditions, different in the southern and northern basins, resulted in enhanced evolution of Caspian molluscs. According to Fedorov (1957), Eberzin (1962), Andrusov (1965), Svitoch (1981), Popov (1983) and others, such guide fossils as the molluscs Didacna Eichwald, which are important for the stratigraphy of Quaternary Caspian marine sequence, are divided into

79

groups of mollusc affinities, namely. D. catillus, D. crassa and D. trigonoides. Each of them includes several lineages, which enabled N.I. Andrusov and P.A. Provoslavlev to recognize five regional stages, Neo-Caspian inclusive (Table 3.2). For the history of their recognition see Fedorov (1957), Markov et al. (1965), and Popov (1983). Table 3.2 presents detailed stratigraphic schemes for the Caspian Pleistocene based on data available as of 1984. The scheme for the Caucasus coast and the northern Caspian area was prepared by the author using data of Alizade and co-workers (1972, 1978, 1984); Zubakov and co-workers (1974); Leontiev and co-workers (1 976); Kozhevnikov and co-workers (1977); Rychagov (1977); Lebedeva (1978); Svitoch and co-workers (1981), and Yanina (1983). The scheme for the northern Caspian area is based on data presented by Goretsky (1966); Shkatova (1977, 1979); Svitoch and co-workers (1981); Popov (1983); Zhidovinov and co-workers (1984). The data reported by Fedorov (1957, 1982) have been used as well; the recent alternative scheme constructed by Fedorov has been edited by Aleksandrova and coworkers (1 984). The mammalian fossil record is an important tool in dating a succession of climatic events over the Caspian area. Occurrences of the Taman fauna described from Azerbaijan by Lebedeva (1978) are dated as Late Apsheronian. Guide fossils of the Tiraspol faunal complex were reported from the Baku and Urundzhik beds (Fedorov, 1957; Lebedeva, 1978). According to more accurate data reported by Popov (1983), occurrences of the Singil fauna including Palaeoloxodon antiguus, are related to the upper Gyurgyan ( = lower Khazar) beds of the lower Volga River. In the northern cis-Caucasus area, along the Kuma river, the Singil fauna was found in travertine mounds of Mount Mashuk at Pyatigorsk, with Th/U age exceeding 300 ka (Cherdyntsev, 1969). The Volga ( = Khazar) mammalian complex, containing Mammathus chosaricus, Camelus knoblochi, Equus chosaricus, Arvicola chosaricus, etc., derived from the Cherny Yar sands yielding Khazar (“upper Khazar”) molluscs (Popov, 1983). The Nikolsk mammoth complex with Coelodonta antiquitatis (Shkatova, 1977, 1979), along with the widely known Binagady complex, comprising some forest faunas, namely the bear Ursus arctos, porcupine Hystrix vinogradovi, deer Cervus elaphus, hedgehog Erinaceous europaeus and the like, which do not now inhabit the desert Apsheron Peninsula (Alizade et al., 1978), is associated with the buried Girkan terrace, 45 - 50 m, and its liman - alluvial equivalents. The age of Caspian climatostratigraphic units has been determined using chronometric methods, such as radiocarbon (Geochronology of the USSR, 1974; Svitoch et al., 1981), thorium-uranium (Arslanov et al., 1978), thermoIuminescence (Leontiev et al., 1976; Svitoch et al., 1981), and the fission track method (Koshkin, 1984; Ganzei, 1984). The data obtained are presented in Table 3.2. A systematic inconsistency between the 14C, Th/U and TL datings should be noted. Thus, the 14C and TL ages of the Lower Khvalyn range from 12- 18 ka to 53 - 71 ka, respectively; and I4C, Th/U and TL ages of the Girkan are in the range of 25.3 -27.6 ka, 76- 81 ka, and 130 + 15 to 143 & 19 ka, respectively. According to Yeremin (1986), in the section along the Shura-Ozen River, Dagestan, Th/U datings of 76 - 8 1 ka were obtained from beds just beneath the Blake excursion (1 13

80

ka) (Third all-Union meeting . . ., 1986, p. 175). This suggests that TL datings best conform to the relative geologic age of the deposits. A K TL age of the Upper Baku-Urundzhik of 400-480 ka (Svitoch et al., 1981) is not inconsistent with the data obtained from fission tracks at 510 f 40 ka (Koshkin, 1984) and 608 +- 70 ka (Ganzei, 1984). In turn, fission track datings of the Lower Baku, at 700 f 200 ka (Koshkin, 1984) and of the Tyurkyan, around 950 - 1,050 ka (Ganzei, 1984), are consistent with paleomagnetic data, namely the Brunhes - Matuyama transition in

Table 3.2. Climatostratigraphic units of the Caspian Pleistocene ~ ~

T C T T Slagc

__c--

-I

~

~-

~

1

-~

cauca\u, coa,t

_

_

North Caspian lowlands

~~~~-~ Tpperl(hvaK- E , l o , r l r s i l e r r a c c i . - I I -6. and - 2 - 0 m. “C I 4 - 30.9 r l ” and r2 rxcur\ionr

Terrace 34 - 36 rn

1

____

1

1st ferrace

rl and r2

Enornerka regrcsrion

excursions

1 1 -

Jc

-1

I 1

Terrace 48- 52 m - Abeskun beds of Manych Stralt, KTL 53 - 64 Binagady iauna. Burlay beds of Manych Strail

Elton regression llnd terrace

Lower Khvalyn

D

ororrucfa

D . ebrrsrnr 5 1 - 71

“6 II - 18. KTI. -

Akhruba regrers,on

Upper Girkan beds of Manych Strait

D.crisfam

41 ;

2

23

‘‘

$ 0 8 l4

Cherny Var beds

D.pulaeofngonotdes beds Upper Kharar MTL 235 -275 r 40

Terraces 100- I05 and 120-130 m D. schuroosenrca. KTL 254 - 340

Terrace 145 - I50 m D. nalivkm

___

Terrace 160- 170 rn Ash A?. F.t 400-602?

k:‘a._.

Louer Kharar

__

Hurcanio grbba

Middle Aprheron Aps. proprnquo. Arch mendrandis

VOlgan fauna Mum. chosurrcus Arv chosorrcus

1

Zaimishche Raigorod beds Singilian fauna Pul. onriqrrus

_______

D. coldlus

devexo

regrercion

- --

I

_

Volga River

Apr. propinquu

Seroglarovka beds Aps. propinquo

,

_.

81

the upper Tyurkyany Formation and the Zykh episode with a K/Ar age of 820 f 250 ka (Zubakov and Kochegura, 1971, 1973) at the top of the Apsheronian. Table 3.2 shows that five traditional Caspian regional stages comprise, all in all, eleven climatostratigraphic units, namely, kryomers, associated with climate humidification and Caspian transgression, and thermomers, corresponding to climate aridization and Caspian regressions. According to Abramova (Fig. 3.3), annual temperature fluctuations reached 7 - 8°C. During the thermomeric and kryomeric phases temperatures were, respectively, 0.5 - 1 "C higher and 4 - 6°C lower than the present. The appearance of the molluscs Corbicula fliiminalis in the Girkan of the northern Caspian area, where they do not live at present, as well as Ca/Mg ratios for valves of brackish-water molluscs of this age from Azerbaijan, made KhalifaZade suggest that in Girkan time the mean water temperature was probably 1.5"C higher than at present (Shkatova, 1977, 1979). The presence of such subtropial elements as Paiaeoloxodon antiquus and Macaca in the Singil fauna implies higher temperatures of the Singil thermochron as well (Lyubin et al., 1985). Throughout the Pleistocene, changes in atmospheric precipitation were much larger in amplitude as compared with the temperature variations. According to Abramova (Fig. 3.3), precipitation was at the present level or decreased to 50- 100 mm during arid periods, and increased to 300 - 400 mm during coolings at transgressive phases of the Caspian. These reconstructions comply with the data available on the history of mountain glaciation in the Caucasus. Milanovsky (1 968) recognized four glaciations in the Greater Caucasus range in Pleistocene time. The threestage Chegem Glaciation was synchronous with peaks of the Apsheronian transgression. According to Milanovsky, the maximal three-stage Eltyube glaciation can be correlated with the Baku transgression, but, according to our correlation, it is related to the Baku and Gyurgyan transgression (Fig. 3.3). The two-stage Terek glaciation is correlative with the Khazar transgression. The last, Bezengi, glaciation is subdivided into seven or eight stages, each probably synchronous with one of seven or eight Khvalyn terraces. The lenght of mountain glaciers reached 50 km at the maximal stages. The trans-Caspian area - the Greater Balkhan Mountains (Turkmenistan) and the Kopet Dag, as well as the mountains of medium height in Iran - was also affected by glaciations. Selivanov (1984) reported that the snowline lowered there to 1.2-0.9 km and the length of glaciers reached 2 km. Very interesting paleoclimatic reconstructions for the Caucasus Range have been made for the sections of Paleolithic cave sites, namely, the Azykh site in Azerbaijan (Velichko et al., 1980) and the Kudaro site in southern Osetia (Lyubin et al., 1985) which represent time intervals unique in duration and covering the Eolithic (Azykh site, bed 15 - 14), Acheulian, and Mousterian. Bed-by-bed collection of bone remains and pollen diagrams served as the bases for the quantitative estimation of temperature and precipitation variations with time. Unfortunately, chronometric datings of the beds in these sites are lacking and allocation of inferred parameters to the stratigraphic scheme is provisional (Fig. 3.3). Velichko and co-workers (1980) reported that in the interval from the pebble culture beds ( = Oldowanian?) to the Early Acheulian beds variations in winter temperature and precipitation reached 9 - 14°C and 500- 1000 mm, respectively. In the interval between the Middle Acheulian ( = Singilian?) and Khazar thermochrons variations in summer

82

temperatures may have been no less than 5 - 9°C and precipitation decreased by one third. Although these estimates are approximate, they are too high for the Azykh situated at an elevation of 440 m and hence need further investigation. However, multiple changes of forest taphocoenoses with Hyslrix and Macaca, and polarsteppe taphocoenoses with Cuon alpinus - Chionomys, and forest zoocoenosis (Velichko et al., 1980; Vereshchagin and Baryshnikov, 1985; Lyubin et al., 1985) throughout the sections of both caves lend additional credence to large-amplitude climatic variations over the Caspian area in Pleistocene times. A similar pattern of changes in vegetation-geographical zones is suggested from spores and pollen analyses performed by A.A. Chiguryaeva for the lower Volga area, E.A. Bludorova for the middle Volga and Kama area, and V.K. Nemkova for the cis-Ural area (Yakhimovich et al., 1984). Resultant geobotanic curves (Fig. 3.4) record up t o 89 climatic rhythms in the Volga Pleistocene, when the boundaries of geographical zones shifted by 1,500 - 2,000 km. Unfortunamely, correlation of these geobotanic phases and fluctuations in Caspian sea level remains uncertain. According to the above authors, forest zones in the deveIopment of vegetation in the Volga area coincide with thermomers and Caspian transgressions, but this is inconYoLga and Kame

ualLeys and UraL Piedmont

Ill 7 1213)4]516171

I-st terrace

Ill-rd terrece

M.chosaricus

Fig. 3.4. Vegetational change inferred from pollen analyses of: key sections of the Northern Caspian area (after Chiguryaeva), Kama River basin (after Bludorova) and Ural Piedmont (after Nemkova). Simplified modification from Yakhimovich et al. (1984). 1 - cold steppe; 2 - light birch forest; 3 - birch-pine forest; 4 - dark coniferous forest (taiga); 5 - coniferous deciduous forest with broad-leaved species; 6 - deciduous forest dominated by broadleaved species; 7 - forest steppe and steppe (xerophytic).

83

sistent with Abramova's data (1982). Hence, the problem of changes in humidification of the Volga - Caspian region and the genesis of Caspian transgressions remains unsolved and highly debatable.

3.3. The Azov-Black Sea basin as a standard for the climatostratigraphic sequence on the shelf off Europe A chronostratigraphic division for the Pleistocene of the Black Sea area was first proposed by Arkhangelsky and Strakhov (1938) and Andrusov (1965) who recognized four regional stages. Later the division was developed by Goretsky (1955, 1982), Fedorov (1963, 1982), Nevesskaya (1965), Zhizhchenko and co-workers (1968), Lebedeva (1972), Ostrovsky (1974, 1977), Chepalyga (1980), Kitovani and coworkers (1980, 1982), Popov (1983) among many others. A Pleistocene deep-sea section was drilled during Leg 42 of Glomar Challenger (Ross, Neprochnov et al., 1978, Zhuze et al., 1980). In 1984 the Interdepartmental Stratigraphic Committee of the USSR adopted a new scheme constructed by a group of investigators (Goretsky, Lebedeva, Shevchenko, Veklich and others) headed by the author. The scheme served as a basis for the present description (Table 3.3). An alternative viewpoint of Fedorov and co-workers was reported by Aleksandrova and co-workers (1984). The Sea of Azov and Black Sea basin is a unique area for climatostratigraphic study of the Pleistocene marine sequence. The basin is linked with the Mediterranean through the Bosporus, a narrow strait whose depth is 36 m at present and it probably has never exceeded 80 - 100 m. Therefore, during glacio-eustatic transgressions the Black Sea was an embayment of the Mediterranean Sea, while in time of regressions it could act as a freshening basin with water flowing into the Mediterranean or as an inland lake. In both cases the level of the Black Sea depended on a river run-off. Hence, the level records climatic changes in the Black Sea drainage area (50- 40"N) and its relationships with the Earth's kryosphere. The Sea of Azov part of the basin is especially sensitive to a regional increase in precipitation due to the influx of the Caspian waters marked by the molluscs Didacna Eichwald into the Sea of Azov through the Manych Strait. New data reported by Zubakov and co-workers (1982) suggest that during the last 1 Ma the Black Sea suffered nine phases of salinity increase accompanied by invasion of the Mediterranean fauna, including six phases recorded in the western Manych (Goretsky, 1982; Popov, 1983). During these phases salinity varied from 15%0 to 27%0. High-salinity phases alternated with nine phases of salinity decrease, six of which were associated with invasion of the Caspian water and fauna into the Sea of Azov. These relations are easily discernible in key sections of the Kerch-Taman type area (Fig. 3.5) and provide a basis for successive climatostratigraphic division of the Black Sea Pleistocene. The division can be correlated with the oxygen-isotope scale of the Pleistocene deep-sea sequence (Table 3.3). Description of the climatostratigraphic division of marine and coastal sequences of the Black Sea basin is reduced here, for simplicity, to four reference sections. Two extend along the coast of the Kerch Strait, a third is on Lake Chockrak (Kerch

84

Peninsula), and a fourth is exposed in the Tsvermagala Mountains in Georgia (Figs 3.5 and 3.6). A possibility to trace facies relations of marine and soil-loessal and alluvial-deltaic units in coastal sections of the Sea of Azov and the Kerch Strait allows for climatostratigraphic reconstructions. During the last 1 Ma u p to 14 buried soils (Fig. 3.5) correlative with marine thermomeric units can be traced in coastal sections. Loessal horizons are equivalent to low-salinity kryomeric units of the marine section.

ELtigen

Lake Tahechik

I1

Th/U 97.3 I2

- 32 - 28 - 24

Cape T u L a

- 20 - 16

x nt

- 12

12

r-

b 16

12

12

8

8

4

4

0

0

Fig. 3.5. Six Pleisrocene key sections of the Kerch-Taman area (height and distance in meters). 1 VI - marine terraces. Legend: 1 loess, 2 - buried soil, 3 - embryonic soil, 4 tongues of redeposited soil, 5 - lagoon silt and clay, 6 - detritus shell-limestone, 7 - dune sands, 8 marine mollusks, 9 - non-marine mollusks, 10 - serpulitic biohersm, 11 - mammal remnants, 12 - frost wedge polygonal structures, 13 - dated patterns, 14 - normal polarity, 15 - geomagnetic excursions. ~

~

~

~

85

Age determinations of the marine and soil-loessal coastal sections of the Black Sea area are based on mammalian and archeological data-planes, magnetic-polarity planes, as well as on different chronometric data, all of which are independent and mutually complementary. It should be noted that the coastal sections of the norTable 3 . 3 . Clirnatostratigraphic units of the Pleistocene of the Black Sea Region ( K M - kryorner, TM - thermomer, age 10' y r ) Suhaerral deposits

Upper loess KTI 4 0 t 5

1

I 'K'

1

2-4 A-

Kegre%iion

lVth ierracc. 30 3 5 m IAiheI - Pophru. T h i U 13'4 ~

Krgres5ion

r

___Vlh irrrace. 3 0 - 5 4 m (Makopse) __x_yI__.-

Regre\<,"" ~

_

_

Vlih terrace, 32 - 68 rr (Pshada) - Puphro

----

RTL 300 * 75 Chelyadm ( = Yalpug) KM. M . primrgenfus TTL > IHO(KT1 500 + IZO?) __ Akrav i - K o n t h r l ? ) T M ~

___x_y_z--

Ncchepsugo KM D. pon/urospro PlonXufo

~-

i

Vllfh terrace 37 - 88 m (Shapsugo) - P o p h f o

71h

~

SO11

,

l3

'

14

N Y y I I I

Ureki K M . K T L 330 D. subpvrmmdoro ~Vlllth terrace, 50-150 m (Idukopas) - C.cf ruhrrculoiurn?

-81h sod

Upper Chauda

Rahu - Niirrui -Plato\o K M D. porvvlu.

RTL 520 t 130

Shava beds (KM?) D. parvulw/orrnrs

~

1

, Elanchik RTL

~

(=

Nisrrw) K M

-I6

KTL. 620 - 640

~

loth

COll

pro1 p<,rrenris

D. buerrcrursu,

18

KTL 600 t 80

Interrupimn

raganrog- Kolkotova T M / M e

m

3 c U

z

e

h

e

d

13th soil

i (TM)

23

~

K\emonatanebt KM

24 -

C h a k h r a l a I M , D. povlovoe Gurian srage - Digressoduocnu digreAsu

~

14th 5011

Sand,

25

3 00

Estimated climate ue

Recent: Ts 2O0G,T, 4"C,ZP8OOmm D.paruulaeformis GeneraL Dreissena ctcaspia composLtron 0 40 80% D.Lschaudae tschnudae

Winter temperature, OG -2-1 0 1 2 J 4 5 I

I

I

I

I

I

I

r l 22

1

wv

-23

'j

24

P

D. D.crassa tschaudae guriensis D. cf:plerstopleura

L

20

19 21

31;

Candona negLecle Theodoxus paLLasi Micromelania caspia

Micromelania so. TrachyLeberis recepta EaccuneLLa dorsoarcuata Lozoconcha parallels L.gibboides

19

-

I

Ts O C

plherissa bogalschoui rach Leherrs pseudconueza

D.tschaudee D. crassa yuriensis D.pLeistopLeurd D.pseudocrnssa D.pLeistopLeura

ion

1st

25

--

' 0

Arboreal pollen (Trees) No arboreal pollen (Herbs)

1

500

I

I

I

700 800 Z 600 presipitation, mm

I

900

26

A @Ores

Miocene

Fig. 3.6. Tsvermagala Mountain key section, Chauda Stage (after Zubakov et al., 1975, with supplements). Pollen analyses performed by Shatilova. Winter and summer temperatures and annual precipitation estimated by Liberman and Muratova. Note two versions of climatostratigraphic correlation of the section.

87

thern Black Sea area are stratotypes for the recognition of the Taman (Gromov, 1948; Lebedeva, 1972), Nogaisk and Odessa (Shevchenko, 1976) faunal assemblages, as well as the Platovo and Khadzhibei faunal subassemblages of the Tiraspol assemblages. Occurrences of the Singil fauna are associated with the Babel ( = middle Uzunlar) beds, whilst the Volga ( = Khazar) fauna is correlated with the upper Uzunlar, Chelyadintsevo and Geroevskoe I beds. Thus, all the mammalian data-planes show a clear relation to the stratigraphy of the Black Sea sediments. Continuous coastal sections, of the northern Black Sea area are ideal for the development of a detailed magnetostratigraphic scale for the Pleistocene. Independent investigations performed by Kochegura and Pisarevsky in collaboration with the author (Geochronology of the USSR, - Zubakov, ed. 1974; Zubakov et al., 1982; Vlasov et al., 1983), as well as studies carried out by Vigilyanskaya and Dudkin (1982) and Tretyak (1983) resulted in the establishment in the Brunhes Epoch of 8 to 11 excursions in a number of parallel sections, stratigraphically allocated to marine sequences. TL datings of one half of the excursions indicate that they are coeval with those recognized in lavas (Champion et al., 1981) and deep-sea sediments (Wollin et al., 1971, 1977). This allows us to consider the Black Sea sequence of excursions as an important tool for chronostratigraphic correlation (Fig. 2.14). Chronometric methods used to date a succession of the Black Sea kryomers and thermomers include, in addition to radiocarbon dating, uranium series (Th/U) datings of molluscan valves (Arslanov et al., 1975, 1983), bone datings from fluorine - chlorine - apatite ratios (Troshkina, in Zubakov, 1974), and TL determinations on quartz. The latter have been carried out at three laboratories, namely the Ukrainian Academy of Sciences in Kiev (V.N. Shelkoplyas, index KTL); the Department of radiochemistry, Moscow University (V.K. Vlasov and O.A. Kulikov, index RTL); and the Estonian Academy of Sciences in Tallin (G.I. Khyutt, index TTL). Table 3.4 presents a comparison of age values for coeval geological formations obtained by different techniques. The table shows that final radiocarbon dates for molluscan valves from the Surozh and Karangat horizons are younger as compared with U-series and T L ages. Therefore, 14C ages for the Surozh transgression at 24-27 ka suggested by Ostrovsky and co-workers (1977) and Chepalyga and COworkers (Stratigraphy and lithology . . ., 1984, pp. 178- 179), and, moreover, the I4C age of the Karangat transgression around 24-27 ka (Semenenko et al., 1976) appear to be methodically unsubstantiated. U-series (Th/U) ages of valves collected from one and the same horizon show a great scatter and inversion of the data obtained and may be regarded as valid only if they coincide with independent values obtained by other methods. Correlation of TL datings with a succession of geological bodies and geomagnetic markers suggests that TL values and accepted ages for excursions are consistent. Nevertheless, there is some discrepancy between the TL age values for one and the same horizon, particularly between those obtained for deposits in different facies. For example, RTL ages obtained for the Malyi K u t and Middle Uzunlar marine beds are 570 and 580 & 140 ka, while RTL ages for the replacing loess-like loams and soils are 500 k 120 and 450 k 100 ka; KTL ages of 285 -t 48 ka and RTL ages of 275 f 70 ka and 300 k 70 ka were obtained for

83

Table 3.4. Comparison of the numerical age of the geologically synchronous horizons obtained by different methods

--

~~~

~~

-

~~

-

_____

~

-

Srratigraphic levels

U - se r i es

Geomagneric

TI.

datum

~-

~~

Surozh

-

-

Dating tcchniquei (age m ka)

Manych: 34.5 Chushka: 56-76

r3 ( - Luchamp, 3Y -43)

Elugen. TTL 52 5

In upper

boundary

~~

Krotkov -

Cape Krotkav:

Burtass

88.9 - 100.5

( = 5d?)

-

Turla: 33 - 66 Eltigen: 90.6 - 125 Maly K u t . 95.2 - 129

( = 5e)

232-42.1

M

E

9

113)

-(

Malyi Kut: R T L 120 i 30

Lake C'hakrai,

Geroevskoe I1 ( = 6)

RTL I65 f 30 Kortyenki. 170 2 40

~

Ceroewkoe I

r 4 ( = Blake.

29.6-39

_ _ _ _ ~ _

~~

Elrigen.

Tobechik ( = 8-9?)

57.9- 70.5

-

~

Eltigen: TTL > 152 R T L 205 t 50 Tsokur. K T L 285 ? 48 TuAa: RTL.: 270 i 60. 300 t 10 Chokrak: RTL. 300 ? 75

rft ( = Levamme,

290 330)

.-__~___-

+

Mamarussian Malokutian ( = PalaeoeuxineI ower Khazarian) ( = 12-14)

100. 530 2 I: Tuzla: RTL- 450 Malyi Kui: RTL: 570 f 130 Chokrak: RTL. 5RO t 140 Tobechik: RTL 5 0 0 f 120 Ureki. K T L > 330 ~

~~

_____

_

_

~

"Baku Mounlam". K T L 480 'Trvermagal. RTL. 520 110

_ _ _ _ _ _ _ _ ~ ~ Brunhes/Matuyarna reversion

Shava: KTL. 600 i 65 T50kur: KTL: m t 70

_ Isokur KTL: ROO f 112

730

_ - YO0 n

- 970

the Geroevskoe I loess, but the marine units of the same horizon yield a RTL age of 205 k 50 ka. An uncritical approach to the first TL datings for the marine units resulted in overestimated old ages for some horizons (Zubakov et al., 1982; Vlasov et al., 1983). Later these estimates were revised (Krasilov et al., 1985). Let us now consider the climatostratigraphic correlation over the area. Table 3.3 shows eight marine terraces developed along the Caucasus coast of the Black Sea; the terraces were studied in detail by Ostrovsky and co-workers (1974, 1977). Most of them are bipartite in structure; brackish-water molluscs, including Caspian emigrants, are abundant in the lower part of the section, while stenohaline fauna with Mediterranean emigrants is typical of the upper part. The coasts of the northern Black Sea area represent a tectonically stable zone. Diachronous marine terraces occur at the same elevation (Fig. 3.5) and, hence, they are difficult if not impossible to distinguish. Many investigators have been led into

error. Thus, Fedorov (1963) and later Chepalyga (Stratigraphy and lithology . . ., 1986) considered the Surozh, Eltigen and Tobechik terrace units of the Kerch Strait, having the same elevation and uniform paleontological characteristics, as a coeval horizon. Similarly, the Patray, Baku and Litvinov units were combined into the coeval Chauda - Baku horizon with no good reason. Recent data required the available stratigraphic scheme of the Black Sea Pleistocene to be revised (Fedorov, 1963, 1978; Aleksandrova et al., 1984). The main problem in the discussion was the age of the Karangat and Chauda horizons. Arguments for their older ages are given below. The extent of the Chauda regional stage (Andrusov, 1965) is determined by vertical distribution of index species Didacna tschaudae Andr. ( = Tschaudia tschaudae Kitov.). The new data suggest that the Chauda includes 6 or 8 climatomers (Table 3.3). The lectostratotype for the Chauda is the Tsvermagala Mountain in Georgia (Fig. 3.6). The lower Chauda - Tsvermagala is incorporated in the Matuyama magnetothem (Zubakov et al., 1975). The author considers that the Chauda regional stage - Didacna tschaudae Andrus. zone - consists of two substages. The lower - Tsvermagala - Chauda, which is incorporated in the Matuyama zone, is equivalent to the Scyphian lateritic pedocomplex and contains mammalian fossils of the Taman assemblage (Lebedeva, 1972; Shevchenko, 1976). Its base coincides with the Kvemonataneby normal magnetic polarity zone (Zubakov et al., 1975), identified with the Cobb Mountain event at 1.1 Ma (Mankinen et al., 1979)*. Hence, the base of the Chauda regional stage is coeval to that of the Sicilian regional stage of the Mediterranean, dated at 1 . I 5 Ma (Ruggieri et al., 1982). Along with the Didacna complex, pebble with attached Bcllunus shells was found by the author in the lower Tsvermagala unit of the parastratotype Chauda section on the Tsvermagala Mountain. Curdium sp., Corbulu gibba, Custrana fragilis shells were recorded in cores and samples collected from the Chauda units underlying the bottom of the Black Sea (Stratigraphy and lithology . . ., 1986). Thus, the first invasion of the Mediterranean fauna into the Black Sea took place in Jaramillo time (OCT 23 or 25). The transition from the Guria to the Chauda is well defined in the sections of western Georgia (USSR) as a time of last major changes of flora, when a polydominant forest with abundant tropical elements and exotic organisms, including, in particular, the fern Dicksonia, Cyatheu, Polypodium, Pteris, etc., which now inhabit jungles of Malaysia (Shatilova and Mchedlishvili, 1980) gave way to a monodominant forest of the present type. The Early Chauda climate in the Black Sea area must have been much warmer and wetter than at present. For example, in the Natanebi River basin, western Georgia, pollen spectra of the Chakhvata thermomer are characteristic of a mesophillic beech - coniferous forest which had up to ten hemlock species, the last tropial ferns, as well as Engelhardria sp., Plutycarya sp., Zelcova serrata, Liquidumbar styracfua, Fatsia aff. juponica, some species of Magnolia, Symplicos, and the like. According to Shatilova and Mchedlishvili (1980), mean January and July temperatures were, respectively, 0.5"C and 5°C higher than at present. Annual precipitation was also higher than at present. Landscapes of the savanna type with

' Or Jaramillo event, 0.9-0.97

Ma, see Fig. 2.6.

90

valley forests of the Mediterranean type and with fauna including southern elephant, ostrich, elasmotherium, and the like were then developed in the northern Black Sea area (Veklich, 1969, 1982). Winters must have been frostless. The Chumbur - Kvemonataneby kryochron, the first essential cooling, caused (i) the disappearance of subtropical elements from the forest of Georgia; (ii) the development of fir forests with only one species of hemlock surviving; (iii) the replacement of the Taman ( = Villafranchian) mammalian fauna by the Tiraspol fauna; and (iv) the formation of the first loessal horizon - Chumbur - along the northern Black Sea coast. The horizon includes a normal magnetic polarity zone with a KTL age of 800 f 112 ka, which may be equivalent either to the Jaramillo event or to that inferred to occur in the time interval of about 830 ka. The outset of Late Chauda time - the Taganrog thermomer - coincides in the northern Black Sea area with the last episode of distribution of savanna-like landscapes, of which up to now there have survived only the upper lateritic (Shirokino?) soils and the Petropavlovka (= Karay-Dubina) mammalian fauna. In Georgia the Late Tsvermagala thermomer comprises the Kolkhida laterites (Tsagareli, 1964) and a floral assemblage of the Viris-Gele section near the village of Khvarbeti (Chochieva, 1975). The composition of the assemblage points to an extremely wide distribution in Georgia of a forest composed of Sequoia, Metasequoia, Taxodium, Cryptomeria, and Glyptostrobus, and, hence, to a high air humidity. The end of Chauda time - the Platovo - Shava kryochron - coincides with the flux of Caspian waters into the Black Sea and the distribution of molluscan assemblage with Didacna parvula - D. rudis up to the latitude of the town of Tuapse. The Shava beds in Georgia comprise D. parvulaeformis Kitov., D. baericrassa omparetica Kitov., D. pontocaspia, and the Baku ostracodes Leptocythere adulata, etc. (Grishanov et al., 1983). The Platovo kryochron was twostage and resulted in the formation of two loessal horizons. The double Yakhno nexcursion (previously referred to as Ureki 11) with an inferred age of 600 ka was recognized in the upper horizon. The TL age of the Shava beds ranges from 520 to 600 ka. According to new data, the Uzunlar regional stage (Arkhangelsky and Strakhov, 1932) consists of marine section 5 climatorners. Its continental equivalent is the Tsokur pedocomplex. It consists of three buried soils and two loess horizons. The Uzunlar thermochron comprises three peaks of salinity increase and three terraces on the Caucasus coast, represented by the “Uzunlar” malakofauna with Paphia, Cardium, Chione, Balanus. The Early Uzunlar thermochron is associated with the recently ascertained Patray invasion of Mediterranean waters into the Black Sea. In stratotype section of Cape Tuzla (Fig. 3.5) occur Patray sands with Cardium edule L, C. edule lamarkii Reeve and Hidrobia ventrosa are underlain by a thick sequence of dense detrital shaly limestine with Viviparus and Baku Didacna and overlain by lagoon silts, RTL 530 f 120 ka. The Patray transgression is characterized by the combination of the Mediterranean emigrants Cardium edule, C. aff. tuberculatum, Scrobicularia plana, Paphia senescens, Nassa reticulatum, Cerithium reticulata, and the like with the Chauda and Baku relics, such as Didacna rudis, D. eulachia, D. pseudocrassa, D. baericrassa and the like. Evidence for a salinity increase at that time was first record-

91

ed by Andrusov in 1898 in the Gallipoli sections (Andrusov, 1965) and by Goretsky (1955) in the Susat beds of the Manych Valley. The above data were confirmed by the author in 1978 at the Tuzla and Maly Kut sections in the Kerch Strait area. Occurrences of the stenohaline fauna in a 150 m terrace of the Tsvermagala Mountain were assigned to this transgression, although Imnadze and co-workers (1975, 1979) assign them to the Karangat, while other authors (Chepalyga, 1980; Mamaladze, 1980) refer them to the Chauda. In 1984 V.V. Yanko and co-workers (Strafigraphy, lithology . . ., 1984, pp. 191 - 192) recorded stenohaline ostracodes and foraminifera, including Neogloboguadrina pachyderma, Globigerina bulloides, Globorotalia hirsuta, G. crassaformis, Globigerinoides ruber and the like in the Patray ( = Karadeniz after Chepalyga) beds. A warm climate of the Patray thermochron is also suggested by the occurrence of Corbiculafliiminalis in the Maly Kut type section and by the eighth succession of lateritic soils, which is the lowest in the Tsokur pedocomplex of the Azov area. The Malyi Kut kryomer was accompanied by the invasion of the Caspian waters into the Black Sea marked with guide molluscs of Gyurgyan ( = Lower Khazar) from the Caspian Sea horizon, such as Didacna subpyramidata, D. pallasi, D. nalivkini, etc. The Malyi Kut kryochron is recorded by the extinction of the last representatives of the Tiraspol mammalian fauna (Mammuthus trogontherii, etc.), dated by fluorine (FCI/P) and TL methods at 400-420 ka and 450 f 100-580 k 140 ka, respectively. The Middle Uzunlar ( = Uzunlar s.str.) is represented by beds with Abra ovafa and Mitylaster lineatus in the stratotype and parastratotype at Lake Tobechik and by beds with Paphia and Balanus in the upper part of terrace VII - Shapsug along the Caucasus coast (Ostrovsky et al., 1977; Imnadze et al., 1979). At the Tskhaltsminda section Chochieva (1 980) traced the last acme in the development of paludal cypress groves and a progressive replacement thereof by alder groves. The appearance of the Singil (= Babel) rodent fauna with Arvicola mosbachensis (Shevchenko, 1976; Aleksandrova et al., 1984) and abundant molluscs Corbicula fluminalis are important indicators of the Middle Uzunlar thermochron. The fourth (in the Pleistocene) salinity increase of the Black Sea waters during the Late Uzunlar (Aksai - Pshad - Konchek) thermomer was the second whose effects were recorded in the western Manych (Goretsky, 1982). According to Markova (Correlation . . ., 1986) the thermomer is marked by a rodent complex of the Volga ( = Khazar) fauna with Arvicola aff. chosaricus. This fauna, including Mammuthus primigenius and Mammuthus chosaricus as well (Lebedeva, 1972), is also typical of three subsequent climatomers, related to the largest salinity decrease of the Black Sea waters. The Karangat (Arkhangelsky and Strakhov, 1932) or “Thyrrenian” (Andrusov, 1965) regional stage includes eight or ten climatomers. It is separated from the Uzunlar in the Azov - Kerch area by the Chelyadintsevo clay sequence of about 60 m thick with freshwater molluscs Limnaeae, Planorbis and the like and occasional Dreissena. In the deep-sea section of the DSDP 380 and 379 holes the Chelyadintsevo horizon contains an assemblage of cold-waer diatoms Stephanodiscus astraea, Cymatopleura solea, Actinocyclus divisus, A . ochotensis, suggesting a decrease in salinity of the Black Sea waters down to 0.5 - 3%0 (Schrader, 1979;

92

Zhuze et al., 1980). A continental equivalent of the Chelyadintsevo and Geroevskoe I beds is the “Middle Loess”, which is the thickest (up to 8 - 10 m) and the most coarse-grained, with a well defined r-zone dated by TL at 285 k 48 ka. The bipartite Chelyadintsevo - Geroevskoe I kryochron is dared by the first appearance in the Azov area o f the Khazar tundra-steppe mammalian fauna with Mammufhus chosaricus and early mammoth (Lebedeva, 1972). The complex and long-term Karangat interval includes four phases of salinity and temperature increase separated by three phases of low-salinity conditions and cooling. All of them are distinctly recorded in the universal1 known Eltigen section along the western Kerch Strait coast (Fig. 3.5). In the loessal section the Karangat thermomer is related to four chernozem steppe soils which form two pedocomplexes: the Bessergenovka ( = Kaidaki) complex, on the surface of which were collected Levallois - Early Mousterian artifacts made of flint (Lebedeva, 1972; Praslov et al., 1983), and the Beglitsa ( = Priluki - Vitachev) complex with typical Late Mousterian artifacts. The Early Karangat - Tobechik t hermomer consists of two subcycles separated by a hydromorphic soil with Helicella, characterized by a poor Mediterranean fauna with Cardium edule L., Abra ovata (Phill.), Mytilaster lineatus (Gm.), Paphia senescens (Cor.), Hydrobia ventrosa (Mont .), etc. The hydromorphic soil and kryoturbated sands (0.1 -0.5 m thick) with shell detritus and a mixed fauna of marine and terrestrial molluscs or dolomitic beach crust occur at the top of the Tobachik beds. There is also a clearly defined polygonal system of frost wedge pseudomorphs (8 to 10 cm wide and 15 to 60 cm deep, polygons are 0.5 to 1 m in diameter). This kryogenic horizon is called the Geroevskoe I kryomer. The age of this bed is estimated by the TL method at 205 k 50 ka (RTL). The second Karangat cycle represented by the Zavetnino beds also consists of two subcycles: marine gravel and sand (3 or 4 m thick) with Ostrea edulis L., Donax truncatulus L., Cerithium vulgatum, Brug., Solen vagina L., Pholas dactilus L. these subcycles are separated by dune sands with Helieella sp. They are overlain by the Geroevskoe I1 kryomer, dune sands with Helicella krynickii and bones of large mammals. This cycle is facially replaced by loess, which in Chokrak section has a magnetic excursion r5a, dated by RTL at 165 30 ka. The third Karangat cycle represented by the Eltigen beds (Zubakov, 1975), or Karangat sensu stricto, unconformably overlaps the first and second cycles. It consists of sands and gravel (5 t o 8 m) with the richest and most stenohalinic fauna of 60 species including Cardium tuberculatum L., Mactra corallina (L.), Donax venustris Poli., Chlamys glabra L. and Mytilus galloprovincialis Kr. At the base of the Eltigen beds typical cyster banks and serpulitic bioherms are developed. A dozen of U-series dates were obtained from the Eltigen by Arslanov and co-workers (1975, 1983) at 88 ka to 129 ka. They are consistent with a RTL date of 120 k 30 ka for the Dinskaya lagoon beds underlain by the Eltigen sands in the Malyi Kut section (Fig. 3.5). The Burtass ( = Dinskaya) kryomer is represented by the marine beds with the Caspian molluscs Didacna cristata. In the Krotkov section, at Taman, it is characterized by a cold spore pollen spectrum with Picea and is dated by the Th/U method at 95- 100.5 ka (Arslanov et al., 1983). The fourth Karangat cycle - the Surozh ( = Tarkankhut, Chushka) beds - is

*

recognized from the occurrence of the Khvalyn index-species molluscs Didacna ebersini (Goretsky, 1982; Popov, 1983). For these beds Arslanov and co-workers (1983) obtained U-series datings in the range of 47 to 62 ka. In the Tuzla section, the Surozh lagoon beds occur at the top of the geomagnetic excursion r3, which is correlated with the Olby - Laschamp event, 36 - 43 ka (Fig. 2.14). These values coincide with a TTL age of 50.2 ka for a buried soil developed on dune sands overlying the Eltigen beds in the Eltigen section. To determine the paleotemperature conditions of the basin Dorofeeva and Bogatina made about 100 Ca/Mg measurements on calcareous molluscs collected by the author from the Eltigen beds with Curdiurn tuberculurttm. The results obtained suggest that in the summer-autumn vegetation period the temperature of Eltigen sea waters was almost the same as at present, but seasonal temperature fluctuations were much lower, i.e. winter - spring water temperatures must have been higher than now. As suggested by loessal horizons and patterned grounds (ofthe fissure-polygonal type) in the Geroevskoe I horizon, as well as by the presence of tundra-steppe faunal fossils (Lebedeva, 1972), during the Karangat kryochrons, namely, during Geroevskoe I, 11 and Burtass, the climate was cold and relatively arid, as least at the end of the phases. This is confirmed by lowering Black Sea level down to 60 - 80 m prior to the formation of terrace I1 containing C. tuberculatum on the Caucasus coast. Ostrovsky and co-workers (1977) regard this regression as “pre-Surozh”, but the author is prone to correlate it with the Geroevskoe kryochron, when the Girkan mollusc Diducnu cristuta penetrated into the Black Sea from the Caspian Sea. However, the coldest and most arid climate occurred in the Black Sea area during the Neo-euxine kryochron which started, according to TL measurements, no later than 40 ka BP. At that time tundra-steppe landscapes inhabited by lemming, polar fox, reindeer and other representatives of the late mammoth fauna, were developed on the northern Black Sea coast. The Sea of Azov dried out and the Black Sea level was 80 m lower than at present (Ostrovsky, 1977). During that time interval of the Pleistocene loess accumulation was the most intense. The available data are consistent with deep-sea record. For instance, in DSDP Hole 380 A, from a depth of 370-50 m, Schrader (1979) and Zhuze and co-workers (1980) found seven short-term peaks of marine diatoms indicating salinity between 10 and 30%0, separated by longer phases of the development of freshwater diatoms with Stephanodiscus ustrueu and the like, suggesting salinity lower than 3%0. The problem under consideration concerning the age of diatom complexes interpreted by the above authors is not discussed here. Hsu and Giovanoli (1980) relate the onset of this interval to the Jaramillo event. If so, the deep-sea drilling data coincide with terrestrial record presented in this paper.

3.4. The Mediterranean - Caspian paleohydrologic system as a record of global and regional climatic changes The Mediterranean, Black and Caspian Seas, together with their river drainage systems, form a vast hydrologic network whose coastlines and biosedimentary cycles

94

record climatic fluctuations throughout the area. Sections 3.1 - 3.3 suggest a single rhythmic pattern reflecting climatic changes typical of the whole of the system. However, it is known that each of the three constituent subsystems may have had its own hydrologic response to unambiguous climatic variations. This probably indicates that regional variations in humidity and drainage are not adequate to global temperature fluctuations. First of all, we shall try to check whether the Mediterranean - Caspian system had a single rhythmic pattern of temperature variations, and, such being the case, whether it corresponds to the oxygen-isotope scale of the deep-sea Pleistocene. Figure 3.7, a synthesis of Tables 3.1, 3.2, and 3.3, shows that glacio-eustatic raised beaches of the western Mediterranean and marine terraces of the Black Sea are equal in number; they reveal the same tendency for a change in fossil composition and yield similar age values. In fact, nine transgressions - salinizations (except for the Holocene) - have been recognized in each basin with the maximal 2nd Strombus-Eltigen, dated at 120- 127 ka, and the Milazzian ( = cycle “U,”)-Patray slightly younger than 600 ka. Both reached the middle of the Manych Strait. The minimal Portuensian salinity increase occurring near the end of the Matuyama Chron, corresponds t o three phases of the Chauda transgression in the Black Sea. Although extremely rare Mediterranean species are known for these phases in the Black Sea, nevertheless they have been found. All the nine transgressions undoubtedly reflect warmer climate conditions than at present. This conclusion has been earlier confirmed by paleontological evidence. Regressive phases, which are also freshening phases, are concomitant in both basins as well. Naturally, in the Black Sea freshening was much greater (euxinization), as compared with the Mediterranean. Maximal regressions (and freshenings) are the Pontinian-Neo-euxinian with an age of 70- f 15 ka, and the Ostian - Chelyadintsevo - Geroevskoe I, occuring about 280 ka BP. As indicated above, variations in the Caspian Sea level are the opposite to those in the Mediterranean and Black Seas. Maximal invasions of the Caspian waters and fauna into the Black Sea took place in Gyurgyan and Baku time dated at 450-500 ka and 700- 600 ka, respectively. Their ages correspond to those of the Nomentanan and Flaminian erosional phases dated at 430 ka and 700 ka, respectively. The appearance of Tiraspol- Galerian mammal fauna related to the end of the post-Jaramillo Ficarazzi - Chumbur kryomer, 0.9 - 0.85 Ma, and its disappearance, related to the Nomentanan - Malyi Kut kryomer, 0.45 - 0.52 Ma are also concomitant in the three basins. The traditional Pliocene - Pleistocene boundary drawn at the base of the Sicilian and Chauda regional stages and aproximately coincident with the Cobb Mountain normal-polarity event, 1.1 Ma, proved to be coeval in the Mediterranean and Black Seas. Hence, new data confirm a single rhythmic pattern of temperature variability within the three basins. Using first appearance datum, age estimates and paleomagnetic datum planes we get the following correlation between the Azov - Mediterranean climatostratigraphic units and isotopic stages (Fig. 3.7). Of particular importance for the correlation is the first appearance datum (FAD) of Emiliania huxleyi, related t o isotopic stage 7 and dated at 220 ka in the Mediterranean (Parisi and Cita, 1982). In the Black Sea it was recorded in piston cores from the reliable Karangat horizon drilled off Turkey (Ross, Neprochnov, et al., 1978).

(OCT

PM

I

Datum plane

Marine terraces and erosional phases Saprope‘s s1 10 Pant i n i a n

-Em. huxley akrne 68-85

Black Sea

Bditerranean Sea

l/-1~-125

cEm.huxleyi

I

Azov

I

Manych

I

I

Neoeuxinian

Caspian Upper Khualyn

-

I

ka 0

:-100

s5 120 -S6 170

-- 200

s7

Ostian -2311

Geroevskoe I ? 205-300

Upper Khazar

Lake beds

TL 235-340

........... uljJ.1LjJJ.L --‘D.naliukini

Chelyadintseuo

L G.oceanica ca. 420 r P;ililnosa

--

-CI

“l11-rjyJJJ-

-- 300

YoLgan fauna Upper/Lower Khazar Singi Lian

40G

5oo

......

subpyramidata Millazzian - Ul

____i Flaminian

...... .......... Lower Baku

680-706

Sera

-- 600

L t . 510

I,-*,..-Shaua-Platovo-Upper Baku

EL. 700

4,oo

_Upper _ _ _Tsverrnagal -_-rocapso (sma!l)akm ca. 930

-Geph

/-I? 0’

trumalul excelsu ca 1150

LGL.

Cusstun ca. 1050

I

4

..........-. ......... Chakhuata?

G u t a n ?=.:

Duzdag regression 1000

.... Apscheronia propinqua ..................................... port-Katon - MiddLe Apsheron c-

1100

1200

Fig. 3.7. Tentative climato?tratigraphic correlation of marine Pleistocene sequences of the Mediterranean, Black and Caspian Seas (for references see the text).

normal polarity, 2 - reversed polarity, 3 - invasion of the Mediterranean mollusk fauna to the Black Sea basin, 4 - invasion of the Caspian mollusk (hiilia 10 thc Am\. - Black Sea baGn, 5 - Caspian guide species i n the Azok and Black Sea sections, 6 - ash, 7 erosional uncomformirp, 8 - hiatus in marine sedimentation. I

~

~

‘

96

The Gephyrocapsa carribeanica zone is associated with isotopic stages 11 and 9, and with the Uzunlar in the Black Sea (DSDP Hole 379, 100 m deep). Some uncertainty is involved with the introduction into correlation of the Alpine terminology in its Mediterranean interpretation (which does not agree with the Central European notion, see Sibrava, 1972). The fact that the Nomentanan erosional phase (“Riss I”, after Blanc, 1957) is correlative with isotopic stage 12 or 14 calls for some attention. Thus, owing to their complex bio-, climato-, and magnetostratigraphic features and to the presence of stratotypes which can be applied to land methods of studies, the Azov - Mediterranean climathems rather than isotopic stages (to say nothing of the Alpine nomenclature) can be used as a tool for inter-regional correlation of shelf and continental deposits, at least in western Eurasia and northern Africa. By no means does this conclusion diminish the role of the oxygen-isotope scale as a global “measure” for climatic changes in Pleistocene time. Let us now discuss how humidification varied within the drainage basins in the Mediterranean Sea- Caspian Lake system. Apparently, an inversion in changes of the Caspian and Black Sea levels, as well as the position of sapropel layers relative to isotopic stages and phases of freshening of the Black Sea waters may be used as a key in this study. Sapropels are black muds, rich in organic CaC03, formed mainly due t o acmes of diatoms which generally do not occur between the sapropels. Sapropels are known only from the eastern Mediterranean; they are designated there S1, S2, S3 . , . (Cita et al., 1977), or A, B, C . . . (Thunell and Williams, 1983) with increasing age. At the base of sapropel layers the composition of microfossils drastically changes: benthic foraminifers, sensitive to the oxygen regime, disappear. The Arcticulina fabulosa (Seq.) - Bulimina aculata Orb. - B. exilis Brady assemblage starts dominating the section. In the upper part of sapropel layers benthic forms become virtually non-existent and then reappear in low numbers several centimeters above the sapropel in different specific composition (Parisi and Cita, 1982). In the sapropels there are also very few species of calcareous planktic foraminifers, although their abundance increases owing to the presence of Neogloboquadrina dutertrei, making up to 70% of the fauna. An isotope shift (both 6 l80 and 6 I3C) reaches colossal values of 5’700 at the lower boundary of the sapropels (Thunell et al., 1983). The foregoing proves a drastic and strong freshening of surface waters during the deposition of sapropels. Hence, they are a result of stagnation - water density stratification into light fresh water at the surface and anoxic heavy saline water at the bottom. The duration of stagnation phases, as estimated by Muerdter and Kennett (1984), varies from 1 - 2 ka to 8 - 9 ka (Fig. 3.8). I t is a fact that the stagnation phases are a response of the basin to hydrologic and climatic changes in Pleistocene time. But we must answer several questions: what causes gave rise to the changes and at what intervals d o the changes recur? Most of the investigators attribute anoxic conditions in the Mediterranean to the discharge through the Bosporus of glacial meltwater which invaded the Caspian and Black Seas as a consequence of melting continental ice sheets. The idea of a great transit spillway of Siberian meltwater from the Verkhoyansk Mountains through the Angara, western Siberian Lake-Sea, Turgai, Aral, Uzboi, Caspian Lake-Sea, Manych, Black Sea to the Bosporus is attractive to many workers (Volkov et al.,

1976; Grosswald, 1983; Thunell and Williams, 1983; among others). Some investigators attribute the lowered salinity to the build-up of the Nile River run-off due to active monsoon circulation (Rossjgnol-Strick, 1985). Others relate the development of anoxic conditions to the change in current pattern through the Straits of Gibraltar. First of all, it should be noted that climatic curves and, particularly, oxygenisotope curves show no distinct regularity in the distribution of sapropel layers (see Parisi and Cita, 1982, fig. 2; Thunell and Williams, 1983, fig. I). Only some of the layers (S1, S 5 , S10, S11) are obviously in line with climatic optima of thermomers. For example, Thunell and Williams (1983) indicate that the base of sapropel D ( = S 5 ) in core TR 172-22 virtually coincides with Termination I1 (a 5%0 shift in 6 I 8 0 ) . In this case we can observe successive maxima in abundance of temperature-representative foraminifer species, such as boreal Neogloboquadrina pachyderma 10 cm below the sapropel, G. inflata 5 cm below, tropical Globigerinoides sacculifer along the base, G. ruber 5 cm above the base. T h u s , the water temperature rises by 3°C in a 15 cm interval for a period of no more than 2 ka. Sapropel layers S6, S12 are also apparently related to maxima in kryomer stages 6 and 12 of the isotope curve (Thunell and Williams, 1983), while S2, S3, S4, S7, and S8 correspond to minor positive to negative oscillations of the isotope curve. This implies that temperature variations proper, including those responsible for changes in monsoon circulation, could hardly cause anoxia. The hypothesis of the glacial meltwater discharge has not been proven even for the Caspian proper. We share the opinion of those who think that stagnation phases in the Mediterranean may have resulted directly from the reopening of a “marine spillway” between the Mediterranean and Black Seas. Using a common model of climatosedimentary cycles, we shall discuss a resultant hydrologic budget (Table 3.5). Climatostratigraphic cycles, composed of a thermomer and a kryomer, can be

Fig. 3.8. Position of sapropels with respect to the oxygen-isotope curve (based on Glohigeritioides Tuber) in f o u r cores through t h e Sicilian Strait. (After Muerdter and Kennert, 1984). (See Table 3.4).

98

Table 3.5. Clirnato-sedimentary cyclicity in the different hydrological systems -~ ~

Kryosphere

Ocean - East Mediterranean

I

Black Sea

Caspm

Rivers valleys

Loess plam

1 1 flTz{ Warm-humid

1

i

Maximum

5 . Marine

regression

4 Warm-humid

Cool-dry

Warm-humid

Hot-dry

phase Warm-humid

1 Forest

soil

further subdivided into stages, keeping in mind certain facies genetic environment. A natural diversity of cycles is primarily controlled by the fact that warm climatic conditions are not coincident in time with humid climatic conditions. Each hydrologic system responds to the lack of coincidence in its own way. Moreover, the same hydrologic systems (rivers, lakes, and the like) operate in a different way at different geographic latitudes, in different vertical belts, and even at different longitudes. In Pleistocene times a key climatostratigraphic cycle is that related to the periodic formations of continental ice sheets and to their melting. The oxygen-isotope scale of deep-sea sediments provides an integral global expression of a cycle. Whatever the duration (100 ka, 40 ka, or 20 ka), the cycle is divided into six stages. The duration of stages is different at high and low latitudes and at different longitudes. The oceanic cycle (changes in ocean level and associated sedimentation) is related to an integral glacial cycle (not mentioning tectonic and geoid deformation effects on ocean level variations), but divided only into two stages. The Black Sea has its own hydrologic cycle composed of six stages owing to the shallowness of the Bosporus. During oceanic regressions, when sea level dropped below the depth of the Bosporus (which is 35 m deep at present), the Black Sea would become a lake whose level would be controlled by the precipitation regime within the drainage basin. During the third stage the Black Sea would become virtually a part of the Caspian basin. The western Mediterranean would have an oceanic hydrologic cycle. In the eastern Mediterranean, the cycle is affected by the Black Sea and for convenience it can be also divided into six stages. The hydrologic cycle of the Caspian, divided into four stages, is synchronous with the oceanic cycle, but operates in reversed phase. Hence, inasmuch as the hydrologic regime of the eastern Mediterranean is related to that

99

of the Black Sea, and the latter becomes a part of the Caspian during one of the phases, and, ultimately, inasmuch as the Caspian sea level is controlled not by variations in the volume of continental ice sheets, but by the precipitation within the drainage basin, all four steps, including the Pontinian - Caspian river drainage system, can be regarded as a complex multicomponent climatohydrologic system. First, it should be noted that there is no evidence supporting the hypothesis advanced by Volkov et al. (1976) and Grosswald (1983) about a “great run-off” of Siberian meltwater into the Black Sea either during the time interval of 16- 10 ka BP (Grosswald, 1983), or at some other time. Variations in the Caspian sea level are typical of any large lake basin situated at middle latitudes and fed by river outflows. For example, they are completely identical to changes in Lake Bolshoi Khirgis-Nur in western Mongolia (Devyatkin, 1981). We shall return to the subject below, and now restrict ourselves by stating a well known fact that high stands of Caspian sea level marked by the Manych run-off coincide with the very beginning of the kryochron part of the climatic cycle. This follows, in particular, from the fact that the Manych run-off is synchronous with cold water sapropel layers in the Mediterranean (S4, S6, S12). For example, cold water sapropel S6 with an age of 164- 172 ka (Muerdter and Kennett, 1984) can be correlated with the Geroevskoe I1 running-water basin of the Black Sea, where we know the occurrence of Didacna cristata - the guide Girkan emigrants from the Caspian. Earlier data (Popov, 1955; Goretsky, 1982) have been supported by new occurrences of the molluscs collected by Chepalyga (Strafigraphy, lithology, 1984, p. 180) from coastal sections of the Caucasus up to the town of Gudauta. The Early Girkan basin sea level had an elevation of 0 + 25 m (Popov, 1984), and those of Geroevskoe and Mediterranean seas had been, respectively, about 10 - 20 m and 40 m below the present sea level (Thunell and Williams, 1985). A one-sided spillway of the Pontinian - Caspian lake basin, which had been fed by excessive river run-off during the first third of the glacial epoch, may have been quite sufficient to freshen surface waters within the whole of the eastern Mediterranean. Thus, the inferences made by Thunell and Williams (1983) show a very good agreement with the data obtained by the author. Variations in the Nile run-off, although fairly significant, did not play a key role. The next “cold and dry” climatic phase can be assigned to the late Early Girkan time, when the Caspian sea level dropped to 15 m below ocean level (Popov, 1983). The Manych outflow ceased and the Black Sea suffered regression and freshening. At that time the Black Sea was a closed lake of the Caspian type. The stages of a closed lake with low-salinity surface layers recurred in the Black Sea hydrologic cycle. This is independently suggested both by diatom analyses of deep-sea sediments (Schrader, 1979; Zhuze et al., 1980), and by the sections of marine terraces along the Caucasus coast of the Black Sea. Ostrovsky and co-workers (1977) reported that in these sections the lower beds with fauna of the Caspian type are usually separated from the upper beds with fauna of the Mediterranean type by a break, which belongs to the phase of a regressive closed lake. Glacio-eustatic ocean transgressions are peaked at a “warm and dry” climatic phase, when the volume of ice on earth is at minimum. During such periods highsalinity surface conditions in the semi-closed Mediterranean reached 38 - 39.6%0. A

100

two-sided connection between the Mediterranean and Black Seas may have been established before the onset of the phase. The surface current originally carried relatively fresh waters (5 to 15%o) through the Bosporus, and the effect of this runoff can be traced in the Mediterranean as far as the Sicilian sill. The deposition of warm water sapropel S5 lasted for 9 ka and spanned a time interval from 125 ka to 116 ka BP (Muerdter and Kennett, 1984). Sapropel layer S5 is coeval with the Eltigen thermomer with a Th/U age of about 124 ka and is no younger that the Blake event at 113 ka BP. At substage 5d, 115 - 105 ka BP, a normal stratification was re-established in the Mediterranean, because the Black Sea was affected by the pre-Surozh regression (Ostrovsky et al., 1977). Cold water sapropel S4, 99- 101 ka old, reflects an Early Khvalyn ( = Abeskun - Dinskaya) discharge of Caspian water through the Manych and Bosporus, as it is synchronous with a cold phase dated by the Th/U method at 100 ka in the section of Krotkov Cape and marked by the appearance of Didacna ebersini in the Sea of Azov. The warm water sapropel S3 was deposited 79- 82 ka BP during a short-term interval. This is well in agreement with evidence on the character of a low Surozh ( = Tarkankhut) increase in salinity in the Black Sea. The occurrence of cold-water thin sapropel S2 with an age of 40 - 50 ka BP is an indirect evidence for a short-term discharge of Black Sea freshened water through the Bosporus at the very beginning of Neo-euxinian time. Hence, the established synchronism between the Mediterranean sapropels and variations in Black and Caspian sea level gives the clue to the reconstruction of river outflow and atmospheric precipitation in the Volga, Dnieper, and Danube basins during well ascertained time intervals. But this problem is in need of further investigation.

RCsumt (1) Variations in Mediterranean and Black Sea level, which are glacio-eustatic in character, reflect changes in ice volume on Earth. Variations in Caspian sea level were controlled by fluctuations in atmospheric circulation and by precipitation over the Russian plain. They are opposite in phase to oceanic transgressions and fluctuations of the Mediterranean and Black Sea level. (2) The deep-sea section of the eastern Mediterranean is an inter-regional stratotype; local clirnatostratigraphic schemes of the whole of the paleohydrologic Mediterranean - Caspian system, together with appropriate river basins (Volga, Nile and others), can and should be connected to the marine sections. (3) Sections of the northern Black Sea area with a remarkable alteration of beds with the Mediterranean, Caspian and “lacustrine” mollusc and micro fossil fauna, and with a distinct facies replacement of marine layers by lagoon - fluvial sequences with mammal fauna are a unique transition stratotype in “ocean - continent” correlation. The sections can be regarded as a shelf parastrafotype of global climatostratigraphic subdivisions - orthoclimathems.

Chapter 4

THE LOESS ASSEMBLAGE OF EURASIA AS AN INDICATOR OF CLIMATIC CHANGES IN THE ARID ZONE

4.1. The loess zone of Europe

In Europe the loess zone extends as a continuous belt from northern France to the Danube valley, reaching its maximum width in the Ukraine and in the basins of the Don and Kuban rivers. The climatic stratigraphic units of loess sequences are more detailed than those of marine and glacial assemblages. However, to identify and correlate soils properly is usually very difficult, because their specific features are rather vague. It is possible to identify reliably only the type of the soil, mostly by the faunistic composition of terrestrial molluscs. According to Loiek (Kukla, 1977), the chernozem soils are characterized by a Striata association of gastropods with Helicopsis striara, Chondrula tridens, whereas the brown forest soils are distinguished by a banatica association with Helicigona banatica. These features are clearly pronounced in central Europe; in eastern Europe, however, they become less and less distinctive (Veklich, 1982). All loess deposits reveal the same type of gastropod composition with predominant Colurnella coturnella, Pupilla loessica, Val/onia tenuilabris and so on. The development of the Pleistocene systems for the loess zone involved a concomitant subdivision of the sequences in the river valleys and interstream areas. I n geological practice loess sequences are usually identified by the alluvial deposits in river terraces dated by mammalian fauna and freshwater molluscs. Veklich and Sirenko (1978) have worked out a different, purely pedostratigraphic technique for dating the loess sequences (Fig. 4.1). Their units are based on the sequence of autochtonous buried soils in the interstream areas, which contains certain stratotypes and is well documented. The alluvial or marine deposits themselves are dated by the soil commencing the sequence of subaerial series in the river or sea terraces. The results obtained by both biostratigraphic and pedostratigraphic methods generally agree with each other. However, the scientists adhering to these methods often have a disagreement of principle, which leads to heated debates when deciding which of the stratigraphic sequences, autochtonous soils (Veklich, 1982) or faunistic complexes (Alexandrova et al., 1982; Velichko et al., 1983), are more reliable. Many local Pleistocene stratigraphic schemes of the loess zone have been developed in Europe. The schemes worked out by Burdier for northern France (in 1970 published in Russian by Moskvitin), by Kukla and Loiek (1977) Fink and Kukla (1977) for Czechoslovakia, by Fink (1974, 1975) for Austria, by Pecsi (1982) for Hungary and by Veklich (1969, 1982) for the Ukraine are most noteworthy among them. In this section of the book we have tried to summarize the available

XI

m A

II

Iu

Fig. 4.1, Srrarjfication of the loess - soil assemblages and river Lerraces in rhe Ukraine and Moldavia (after Veklich in Zubakov, 1974, fig. 7). 1 - IX - river terraces; 1 - 7 - soils: ( I ) chernozem, (2) brown and light brown, (3) grey brown, (4) reddish brown, (5) red brown, 16) meadow hydromorphous, (7) embryonic; 8 - loess; 9 - c1ay;lO- 11 - Pliocene beds: 10 - light red brown clay and soil, I 1 - red cinnamonic clay and soil; 12 - sand and gravel. For indexes of stratieranhic units FCC Fiec. 4 7 a n d 1 1 d

103

information mostly in relation to the European part of the USSR, Hungary and Czechoslovakia. In Table 4.1, the column for river valleys summarizes the data on the Dniester and Prut rivers as well as the lower reaches of the Danube after Bukatchuk (1983), Table 4.1. Climatostratigraphic units of the Europe loess area Southern part of the Ruician plain

I

River valleys

2 i

I

Lower level

L o e s ~- ~ o i lsequence

I

1

I tl

Upper (Prichernomorye) loess KTl21-35 Upper

9.

-

Tapiosuly I TL 21 - 2 8

24 - 30 I 3 TL 35 4

TTL > 26

I

IInd terrace

loess 5 Tl45-77 Baraharc A T I 81 t 1 0

6,

Lower level (Korman)

-J

cI_

lllrd terrace (Dukhovskaya Reni)

Upper level Karagash beds PO/ onoquur, RTL 140- 142. Th/U 116 k 2

1

De.sna K M

Mende - Bav s

___

1

Speya T M

1

RTL 235 - 260 y;

Fluvioglacial beds

1.6 (r Blake) TL 87 - 98 -soil MB

Bt

1

Kaydak - Bessergenovka-Kursk Korrhevo s Late Acheuhan, K T I 167-240

- 246

I

Dnicperian - Tsna loess and till KTL 230-285

- Potvaaailovka 2 1 Bokovo Late Acheulian-

-

Rornnv

1

i

F.

'

D -

Orel - Orchik - Morsha I. KTI 322 RTL 340 ? 85 380 f 95

terrace = Streha 14

l5 16

I'

3

_i 2 I 21

F

~ t r e ~ i ts aChiglrin

7

Vlth (Vth?) terrace (Gunki - KolkorovayaNagornoel and alluvium of buried valleys. Tirarpolian fauna

terrace Shamin) R polarity

Kairy)

r

, TL 620

2

Loess I V TL 379 i 56-

m

e

-

Khadrhibey K,W KTL 510 i 70 Uzmari T M FAD Arvicolo so. "sirus ~

- 279

P","S3

y

e" -m

13

1

Loess 111

Zaradorka- Kamensk so11 Middle Acheulian. Arv. chosorlcus

c

KM ~ -

640

Kolkotova TM

lKarai-Dubina 1x1 RTL 670 - 900

--__.__

~

~

~

~

~

,

~

~

~

~

u

J

I04

Bukatchuk et al. (1983), Konstantinova (1967) and Alexandrova et al. (1982, 1984), on the Dnieper River after Veklich (1969, 1982), Veklich et al. (1967 - 1984), Goretsky (1979) and Shelkoplyas et al. (1986) and on the Don river after Lebedeva (1972), Goretsky (1982), Vasiliev (1984), Popov and Rodzyanko (1947), Krasnenkov et al. (1983) and Velichko et al. (1980, 1982). For interstream areas, the data have been derived from Veklich et al. (1982) and Velichko et al. (1980, 1982). There has also been archaeological evidence from the Molodovo sites in the Dniester basin (Ivanova, 1977, 1986) and the Kostyenki sites in the Don basin (Praslov et al., 1983; Lazukov, 1981). The magnetostratigraphic scheme for Carpathian Ruthenia has been based on the data derived from Adamenko et al. (1981), for the southern Ukraine, from Tretyak (1983), Vigilyanskaya and Dudkin (1983) and for the Don river, from Glushkov (Krasnenkov et al., 1983) as well as from Pisarevsky (1983) and the author. In constructing the table we have taken account of the results of the Joint Stratigraphic Quaternary Meeting held in 1984 under the supervision of 1.1. Krasnov and the 4th All-Union Meeting on the Quaternary held in Kishinev in 1986 (The Anthropogene and Palaeolithic . . ., 1986; Correlation . . ., 1986). Among the enormous bulk of information on western Europe, Kukla's scheme (1977) attracts particular attention. It is based on climatic sedimentary cycles (soil - loess) and appears t o be the most comprehensive for the loess assemblage. The shortcoming of this scheme is that it contains neither stratotypes nor numerical dates, except for the Brunhes - Matuyama boundary in the soil of glacial cycle J traditionally called the Gottweig. In addition to Kukla's scheme, the table contains Pecsi's scheme (1982) for Hungary, the units of which have recently been dated by thermoluminescence technique (Maruszak et al., 1983). The table for the European loess zone has been presented in such a general form for the first time. Over the last 1.1. Ma (the Holocene excluded) we can distinguish in the European loess assemblage about 22 - 26 soil stratigraphic units (loess members and soils) and 8 to 10 river terraces. The stratigraphic age of the units has been found by their correlation with the Pleistocene Black Sea sequences, with the help of palaeomagnetic marks and thermoluminescence technique, about 100 dates being obtained by Shelkoplyas et al. (1986) in the Research Laboratory of the Ukrainian Academy of Sciences, Kiev, by Butrum (Maruszak et al., 1983) in Lyublin and Kulikov at the Moscow State University (The Anthropogene and the Palaeolithic . . ., 1986). The regional correlation has been based on: (a) facies replacement and redepositing of sediments from the high second, fourth and fifth fluvial terraces above the flood-plain in the Black Sea basin to marine sediments of the Eltigen, Uzunlar sea level, and Patrai transgressions; and (b) facies interactions of soils, loess and alluvial deposits and two glacial horizons. The upper glacial horizon wedges as a tongue into the loess zone along the Dnieper Valley up to the city of Dnepropetrovsk (48"30'N. Lat.). The lower one enters the loess zone by two tongues, along the Don up t o 49'30". Lat. and near Lvov (50'N. Lat.). The upper moraine in central Europe corresponds to the Saale; the age of the lower one is difficult to determine. Polish scientists (Lindner, 1982, 1984) date it by the San glaciation, Czech scientists (Sibrava, 1972), by Elster I and Soviet scientists either by the Don glaciation (Krasnenkov et al., 1983) or the Krukenichi glaciation (Shelkoplys et al., 1986).

I05

These moraines are sometimes divided by a soil - loess sequence, including four related reddish-brown soils (the Tambov - Gorod pedocomplex of the Don basin). The pre-Don (= Saval) subaerial deposits of the Lower Pleistocene are often separated from the Pliocene by erosional unconformity, which in the Alpine foothills is referred by Penk to the beginning of the Diluvium (see Fink and Kukla, 1977). In the Krems section near Linz this unconformity is represented by pebble gravel (Fink, 1974) overlain with the so-called “Mittlere wand” with Columella colurnella and Vallonia tenuifabris. The underlying soil layers K R 7, 8, 9 of the Krems soil complex of Gotzinger with the fauna of Helicigona banatica and pollen of Celtis are considered to represent the Pleistocene glacial cycles K - L - M (Fink and Kukla, 1977). HoraEek (1977), however, thinks that because of the presence of the remains of Mimomys pliocaenicus and M.reidi, these soils correspond to the Pliocene and the N-zone in the upper soil layers KR 7 to the Olduvai. It is quite probable that HoraEek is right. In the Crimean foothills around the mouth of the Belbek River such an erosional unconformity separates the Pleistocene terraces from the Taurida molasse terminated by the Berezan conglomerates and the Kryzhanovka soil, which is reliably dated by the Pliocene (Veklich, 1982, fig. 52). In the Crirnea as well as in the Alpine and Carpathian foothills the erosional unconformity is palaeomagnetically dated at 1.2 Ma to 0.9 Ma. Thus, it is synchronous with the base of the Sicilian stage and corresponds to isotope stages from 28 - 26 to stage 22. The relevant deposits corresponding to this unconformity (the Ilyichevsk and Lower Shirokino-Urzuf horizons in the Ukraine and Dunafoldvar 1 - 2 in Hungary) contain the Tamanian fauna of mammals ( = Nagiharsonihegi in Hungary) with Archidiskodon meridionalis tarnanensis, Prolagurus pannonicus, Allophaiornys pliocaenicus and so on. In the Dniester valley this transient horizon with A . tamanensis is represented by alluvial deposits in the eighth Kitskan terrace (Bukatchuk et al., 1983). According to Chepalyga and Kulikov (Geochronology of the USSR . . ., 1985), there has been found the Jaramillio Normal Event with RTL date of 940 ? 20 ka. During Shirokino time the Ukrainian climate was similar to the unstable humid climate of the subtropical zone with mild winters without frost and hot summers (Fig. 4.2). However, the valley forests and open woodland did not contain any subtropical relict elements, which became extinct in Ilyichevsk time. The end of the transient stage corresponds to the first loess horizon of the European part of the USSR, i.e. the Chumbur ( = Trostnyanka?). It is associated with the revealed remnants of the animals of the Karai-Dubina ( = the early Tiraspol) faunistic complex with Prolagurus pannonicus - Microtinae gen. (Markova, 1982, Shevchenko, 1976) and Bison cJ schoetensacki (Lebedeva, 1972). The presence of the normal polarity zone and KTL dates of 923 _t 165 ka and 850 t 100 ka allow us to correlate the most ancient loess with the Chumbur horizon and isotope stage 22 or stages 22 and 24. According to Penk’s data given in Fink and Kukla (1977), it is just this kryochron that should be considered as the “Giinz”. There are however other data for comparison, for instance of Voznyachuk (1981) and Sibrava (1981). The upper Shirokino - Balashov soil complex terminates the Scythian series of red clays in the southern part of the European USSR. It is synchronous with the alluvial deposits of the seventh river terrace (the Mikhailovka - Nagibino - Kopan terrace), which has no less than two superimposed suites in the Don basin (Kholmovoy et al.,

106

1985) or two levels in the Dniester valley (Bukatchuk et al., 1981). The lower layers of the alluvial deposits of the Mikhailovka terrace with Crassiana crassoides and Mammuthus frogontherii have reserved polarity and RTL dates of 900 k 200 ka (The Anthropogene and the Palaeoiithic . . ., 1986). The upper alluvium is normal and dated at 675 f 170 ka by the same method. The Ilyinka - Sokal - Saval soil is correlated with the soil of cycle J in central Europe (Gottweig - Krems 4), containing Helicigona banatica and the remnants of Hystrix (see Loiek's data in Kukla, 1977). By palaeomagnetic and TL dating the Upper Shirokino thermochron (cycles J - I) and the lower layers of the Mikhailovka terrace are correlated with the Taganrog horizon and isotope stages 21 and 19. In Carpathian Ruthenia it is marked by the early Acheulian beds with KTL dates of 850 k 110 ka (Adamenko et al., 1981). The spores and pollen analysis indicates that during this time the Ukrainian environment was represented by forest steppe and coniferous/broad-leaved mixed forest with such Balkan exotic species as Juglans and others. At that time the Ukrainian climate was warmer and wetter than at present. The most important chronostratigraphic indicator, i.e. the Brunhes - Matuyama boundary, separates the Ilyinka soil from the above-lying Don - Priazov loess. The

I

Unit

Kaidak

Dniepr 1

Holocene Pncherno

Potyagailovka Ore1 Zauadouka

morski

TiliguL

Doflnouka

Lubny

Sula

l 4tacheu

Martonosha

Udai

PrLazovski

PrrlukL

Shirokino

Tyasrnin

Il'itheusk

Fig. 4.2. Climatic curve (summer temperature) along the Ukrainian soil -loess sequence. (After Veklich et al., 1984).

107

latter is double (or triple). In the Don basin, it probably corresponds to two moraines, one of which, according to Liberman, the lower Lipetsk, is not positively ascertained (The Margin Formations . , ., 1985, pp. 150- 151) and the other, the Don - Kanev - Krukenichi is quite reliable. Until recently the latter moraine has been referred to as the Dnieper moraine in the Don basin and the Oka moraine in Carpathian Ruthenia. Actually its age is no less than 510-520 ka by recent KTL dating of the Kanev moraine carried out by Shelkoplyas et al. (1986). The palaeomagnetic evidence has also shown that the Don moraine has two excursions, which are correlated with the excursions of Yachno I and I1 (Table 2.4). The Don - Priazov horizon is defined by a typical middle Tiraspol fauna ( = Nistrus Platovo - Klepki - Bogdanovka - Templomhegi) with Bison schoetensucki lagenocurvis and Prolagurus posterius. The Martonosha thermochron, which is reliably identified by a soil complex of the same name in the interstream sections, is correlated with the alluvial deposits of the sixth Kolkotovo terrace of the Dniester. The age of this thermochron is determined by facies replacement of the alluvial deposits of the Kolkotovo terraces containing Viviparus tiraspolitanus in the lower reaches of the Prut and Danube by the Uzmari estuary-marine layers with the Late Baku mollusc fauna: Diducna pseudocrassa, D.exgr. tschuudae (The Anthropogene and Palaeolithic . . ., 1986). This fauna indicates that the Uzmari layers refer to the Patrai transgression (the 15th orthoclimathem). The mammalian fauna of the Kolkotovka terrace is represented by the Tiraspol complex and is similar to the Sussenborn and Mosbach fauna. The former generally comprises Mummuthus trogontherii, Bison schoetensacki schoetensacki and Mimomys infermedius (Alexandrova et al., 1984). The Sula kryochron is represented by a loess horizon of the same name (Veklich, 1968). In the Don basin it seems to correspond to the Korostelev loess after Velichko et al.’s system (1982). This loess and the upper fluvial deposits of the 6th terrace are marked by the last finds of the animal remnants of the Late Tiraspol complex ( = Khadzhibei) with M. trogontherii and the arctic-steppe migrants Dicrostonyx sp. and Lemnus sp. According to Kalke, some of the elephant’s teeth found in the Sussenborn and Upper Kolkotovo layers refer to the archaic rnamrnuth, Mumm uthus primigenius. The TL dating of the Martonosha - Sula climato-sedimentary cycles and its Hungarian equivalent is ambiguous. Shelkoplyas et al. (1986) date the Khadzhibei sands of the Gunki section at 510 k 70 ka. Vlasov and Kulikov obtained much greater dates for the Kolkotovo section: from 620 - 630 k 150 ka for its upper part to 735 - 800 ? 160 ka for the lower part (The Anthropogene and Palaeolithic . . ., 1986). Butrum dates the Paks-Dunakolod soil at 422 k 64 ka (MaruszLk et al., 1983). After Troshkina (Geochronoiogy of the USSR . . ., Zubakov, ed. 1974), the fluorine method gives the age of the Late Tiraspol fauna as 402 ka. The IVth (Vth?) terrace above flood-plain of the rivers of the Black Sea basin (the Mariinsk terrace on the Don, the Chigirin terrace on the Dnieper, the Grigoriopol terrace on the Dniester and the Dzhurdzhuleshti terrace in the Prut valley) corresponds to two climato-sedimentary cycles and consequently, to four climatomers (Table 4.1). The first cycle is represented by a deeply wedged alluvial bed with the Chigirin ( = Singil) forest fauna (in Europe it is the Mauer fauna), in which

108

Palaeoloxodon antiquus and Arvicola mosbachensis appear for the first time (Markova, 1982). According to Bukatchuk et al. (1981), in the lower reaches of the Prut and Danube the alluvial deposits of the Vth (IVth) terrace are facies replaced by estuary-marine Babel (“Euxine-Uzunlar”) layers with Didacna pallasi, D. nalivkini and warm-water Corbicula flurninalis. The TL age of the alluvial deposits of the IVth Dniester terrace near the village of Varnitsa is 340 k 95 ka (The Anthropogene and Palaeolithic . . ., 1986). The KTL dates of the Lubny soil are 410-440 ka. The climate of the Lubny - Inzhavino thermochron was warmer and wetter than the present climatic conditions. This inference has been made proceeding from the nature of the soil, palynological evidence and water and land mollusc assemblages that included Gastrocopta theeli and Corbiculaflurninalis which occur at present in the Caucasus. Gubonina has identified in the sapropel of the IVth Dnieper terrace that is close to the village of Gunki a spectrum of coniferous and broad-leaved mixed forest with patches of firs, spruces, hornbeams, lindens, oaks, elms and alders. None of the exotic species have been identified. Judging by the presence of Carpinus betulus, which grows at present only in Carpathian Ruthenia, the mean temperature in January is supposed to be -6°C and in July, + 18°C and the total annual precipitation, about 500 to 600 mm. That means that the climate at that time was milder and wetter compared even with the Holocene Optimum (Velichko et al., 1980). This agrees with the data of Grichuk and Gurtovaya (1981) on the Krukenichi section near the city of Lvov, where only the sediments of the first half of the interglacial containing the evidence of the coniferous and broad-leaved forests with hornbeams (24%), oaks, elms as well as beeches, walnuts, yews and vines have been observed. The Tiligul ( = Borisoglebsk) loess has been dated by TL technique at 400 - 440 ka. The next, Zavadovka - Dnieper climatic cycle corresponds to the upper part of the section of the IVth terrace. It is defined by the Volga - Khazar fauna with Arvicola chosaricus (Markova, 1982) and the Middle AcheuIian artifacts. The KTL date of the Zavadovka soil is 330 t o 380 ka. The brown, Potyagailovka (= Romny) soil is often inseparable from the Zavadovka soil. The Dnieper ( = Tsna) loess is facies replaced by the fluvio-glacial sands and tills. The palynological evidence shows that the entire southern part of the European USSR was in periglacial cerealwormwood (Grarnineae, Artemisia, Chenopodiaceae) steppe environment with sparsely growing pines in the watersheds. Eternal frost spread southward up to the Black Sea, which then experienced a regressive stage. The loess sequences exhibit the boreal - alpine assemblage of land molluscs with Vallonia tenuilabris and Columella cofumella among others (Veklich et al., 1977). A series of KTL dates for the Dnieper loess and till is within the range of 240 to 285 ka (Shelkoplyas et al., 1986). The TL dates and the Volga complex of mammals explicitly show that the Dnieper kryochron was synchronous with the 8th orthoclimathem. The Kaydak - Tyasmin climatic rhythm corresponding to cycle C of central Europe is sometimes still omitted by a number of scientists. Its thermomer portion is correlated with fluvial deposits of the IIIrd Dniester and Don terraces above the flood-plain RTL dated at 230 - 260 ka and in interstream areas, with the fifth buried soil counting from the top, which contains the artifacts of the Late Acheulian

Palaeolithic (Adamenko et al., 1981; Praslov et al., 1983). The KTL dates of 180 j, 30 ka and 190 f 32 ka of the Kaydak soil are most convincing for an extensive continuous sequence near the village of Blizhnii Khutora in Moldavia (Bukatchuk et al., 1983). The Kaydak thermochron had two optima: the first gave rise to the development in the Ukraine of grey soils in the fir and oak forest environment and the second of the chernozem and t u r f podsols in the steppe and pine and broadleaved forest environment (Veklich et al., 1984). The Tyasmin kryochron produced a thin loess sequence containing the Levalloisian - Mousterian artifacts KTL dated at between 140 and 170 ka. Vozgrin’s data (The Margin Formations . . ., 1985, p. 115) show that in the woodlands near Zhitomir, loess is replaced by a glacial horizon, which was probably left behind by the Moscow glacier. The Tyasmin loess is comparable with the first Middle Pleistocene loess in Hungary, the “Riss” loess according to Pecsi (1982). The Priluki - Beglitsa - Salyn thermochron is a time equivalent of the Eltigen transgression, as it follows from recent data obtained by Zubakov.’ It is associated with the invasion of the central European fauna (the Banatica complex of gastropods, Corbicula fluminalis, Unio,Pulaeoloxodon antiquus germanicus, etc.) up to the Don basin. The Priluki - Beglitsa brown soils marked by the presence of flinty tools and weapons of the Early Mousterian (Lebedeva, 1972; Adamenko et al., 1981) were formed, according to pollen and spore evidence, in the oak - linden - hornbeam forest environment in the northern Ukraine and in various grass steppe environments in the southern Ukraine. The first half of the Priluki interglacial was warmer and wetter, while the second was warmer and drier than the present time (Fig. 4.1). The age of the Priluki thermochron and its analogues (Salyn, Gorokhov, Mende - Base and others) has been unanimously determined by two independent methods: according to the U-series for the shells from the fluvial deposits o f the IInd Dniester terrace, it is dated at 116 f 2 ka (Cherdyntsev, 1969), the RTL dates of the same terrace are 140- 142 f 40 ka (Geochronology of the USSR, 1985, p. 104 and the KTL dates for the Priluki soil are 100 to 130 ka (Shelkoplyas et al., 1986). In the loess zone we observe an exceptionally full dissection of the sediments of the last glacial periods that form the Antian horizon in the Ukraine. Its stratotype is the upper part of the sections of the second terraces above flood-plain of the Dniester near the villages of Molodovo and Korman studied by Ivanova (1977, 1986 and of the Don near the village of Kostyenki studied for a long time by geologists and archaeologists (Praslov and Rogachev, 1982). According to Veklich (1982), in the interstream section, five climatic stratigraphic divisions can be considered as equivalent of the Antian horizon. lvanova has identified six buried soils and seven loess layers in the sections near the villages of Korman and Molodovo. Similar units are evidently identifiable near the village of Kostyenki. Climatic stratigraphic units of the Antian horizon agree with the history of the development of the Palaeolithic stone artifacts, the finds of which serve as reference indicators. However, the I4C and U-series techniques give in the main much younger dates for the cultural

‘

At one time the authors (Zubakov and Borzenkova, 1983) erroneously correlated this thermochron with the Surozh transgression. Such correlation is now adopted by Veklich et al. (1984).

110

material, which disagree with TL quartz dating. Therefore, in constructing Table 4.1 we have used only the maximum I4C dates, considering palaeomagnetic and TL data to be more preferable. There are identified altogether four polarity excursions in the section of the Antian horizon both in Kostyenki (after Pisarevsky and Zubakov) and in Molodovo (after Dudkin, oral report). The youngest excursion r l ( = Gotenburg, dated at between 12 and 12.6 ka) has been found in Kostyenki at the foot of the present soil and the second, excursion r2 ( = Mono, dated at 28 to 30 ka), which has been dated at more than 22.27 ka by the 14C technique and at 26.76 ka by TTL, is found under the Gmelin palaeosol of the Don first terrace above the flood-plain and under the Bryansk soil on the Dniester (Kulikova, 1980). It appears thus that the age of the Bryansk soil could not be more than 28 to 32 ka. The third excursion r3 ( = Laschamp - Kargopolovo, 41 - 43 ka) has been observed in the roof of the lower Korman ( = Dofinovka) soil dated at 44.4 ? 2.05 ka by the 14C technique (Grn 6807). It is in this soil that the traces of the final flake Mousterian have been found (Ivanova, 1977, 1986). Near Kostyenki this soil (the “Lower humus”) contains vestiges of the sites of the Spitsyn type, i.e. the earliest Upper Palaeolithic exhibiting traditions of the Mousterian culture. Thus, the layer of basalt tuff from the loess separating the upper and lower Kostyenki humus can probably be related to the most widely distributed Mediterranean ash horizon, Ischia 5, which is 38 ka by the K/Ar technique (Thunell et al., 1979, Praslov and Rogachev, 1982). The fourth excursion r4 ( = the Blake, 113 ka) has been revealed in Kostyenki within a thick black soil layer at the foot of a “Stratigraphic hole”, which is underlain with deluvium loam with an RTL date of 170 ? 40 ka (the author’s personal information). It has also been found at the boundary of the Vitachev soil and the Uday loess at the Korolevo site in Carpathian Ruthenia (Adamenko et al., 1981). All these data show that the Uday loess containing the “developed Mousterian” artifacts can be correlated with the Burtas ( = Krotkov) kryochron in the Black Sea area (95 - 105 ka by U series) and isotope substage 5d. The double Vitachev ( = Molodovo, Basaharc A) soil with the Late Mousterian artifacts (Molodovo I and 11) thus seems to be a time equivalent of the Surozh ( = Tarkankhut) transgression and isotope substages 5c - 5a dated at 105 to 89 ka. However, the T L age of the Vitachev and Basaharc A soils does not exceed 80-85 ka. The climate of Vitachev time was dry and warm resembling the present climate. This is well confirmed by the type of soil (chernozem and leached turf-carbonate soils) and palynological evidence (steppe with valley pine forests with extensively distributed broad-leaved trees, including the hornbeam, linden and oak) (Levkovskaya, 1977; Veklich, 1982). This time period is characterized by frequent climatic fluctuations (Fig. 4.2). The Bug ( = Khotylevo) kryochron that left behind loess with the final Mousterian artifacts formed under the conditions of a more severe periglacial climate as compared with the Uday loess. All the Ukraine was then in the tundra steppe zone (Veklich, 1982; Gerasimov and Velichko, 1982). The early Dophinovka ( = Basaharc” D” = Korman I) soil and the top part of the lower humus layer in

Ill

Kostyenki that corresponds to this soil developed during the double interstadial between 55 - 60 ka and 40 - 42 ka. This soil is related to isotope substage 3c and is marked by the transient Middle - Upper Palaeolithic flinty tools and weapons. In Poland this interstadial is called the “Aurignacian” and in western Europe it probably corresponds to the Moershooft and Hengelo warm stages (Van der Hammen et al., 1967; Leroi and Gourhan, 1968, 1977). According to Levkovskaya (1977), Malyasova and Spiridonova (oral report), the upper portion of the lower humus layer in Kostyenki (51”N Lat.) formed in the environment of fir taiga with few broad-leaved species. The climate was cooler and more continental than at present. The chernozem steppe of the West Siberia type was then developed in the Ukraine (Veklich, 1982). The inter-Dophinovka kryochron corresponding to isotope substage 3c (40 - 33 ka) exhibited a severe periglacial climate. The Late Dophinovka thermochron (33 - 24 ka) is represented in Kostyenki by two thin soil layers (the “Upper humus”). The lower soil layer is dated by the I4C technique at 32.7 ? 0.7 ka and the upper one, at 26 - 28 ka (Praslov and Rogachev, 1982). The soil complex Mende “F” in Hungary has the same TL dates of 33.5 ka (PeEsi, 1982). Over the entire southern part of the European USSR, the upper Dophinovka soil is represented only by brown soils. In Kostyenki there are found in this soil the taiga - steppe mosaic spore and pollen spectra, which are indicative of frequent climatic fluctuations with alternating periglacial and boreal environments. The latter is in particular confirmed by charcoal from the Late Palaeolithic fires (made of broad-leaved species such as the oak, linden and hawthorn), which are associated with the cultural horizons of the middle Kostyenki Palaeolithic (Lazukov, 1981). The upper loess layer of the Antian horizon, i.e. the Black Sea loess in the Ukraine (“Prichenomorsky”), the Sungir, in the Russian Plain (Velichko, 1980) and the Tapiosuli in Hungary (PeEsi, 1982), covers not only the second terraces above the flood-plain but a high level of the first terraces. It is associated with the sites of the Late Palaeolithic hunters for mammoths. The lower portion of the loess is dated by radiocarbon at 20-23 ka. It is separated by several humus interbeds, the thickest one (Trubchevski, after Velichko et al., 1980) is distinguished everywhere and dated by the 14C method at 16- 17 ka. The upper part of the first terrace above the flood-plain in the Don, Dniester, Desna and other rivers’ basins exhibits well pronounced thin layers of young buried soils of 12.3 to 12.0 ka and 11.8 to 11.7 ka old. The climate of the kryochron in the Black Sea area experienced, thus, frequent fluctuations and was most severe between 23 and 18 ka ago. It can be seen from Table 4.1 and Fig. 4.3 that the loess sequence in central Europe as it is found in Czechoslovakia and Austria (Kukla, 1977; Fink and Kukla, 1977; Fink, 1975; Kretzoi and PeEsi, 1979) and in Hungary (PeEsi, 1976; Pevzner and PeEsi, 1980; PeEsi, 1982) resembles very much that of eastern Europe. They both have the same number of stratigraphic units, identical palaeontological features (HoraEek, 1977) and the same superhierarchy of units determined by the amplitude and duration of climatic fluctuations. Thus, the Tiraspol loess sequences in the USSR are correlated with the Lower Paks loess in Hungary and the Middle

.

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Fig. 4.3. Environmental changes around Brno and Krems as reconstructed from the loess record (after Kukla, 1977, fig. 17). Symbols of snail assemblages explained in the legend. In the local fauna column, crosses mark single faunal occurrences, full dots first faunal occurrences in the loess sections. Cross-hatched - warm savannah type environment favoring development of exceptionally red polygenetic soil. Breaks are levels of deep erosional incisions.

8

12

1550~ 0

/

20 0

40

\

60

Fig. 4.4. Correlation of key sections of the loess assemblage of the Tadjik Depression (after Lazarenko, 1976; Minina and Lazarenko, 1984). Explanation: 1 - recent soil, 2 - horizons of buried soil, 3 - fully developed zonal-type soil, 4 - moderately developed soil, 5 - slightly developed and embryonic soil, 6 - visible indication of soil formation in the loess, 7 - clay, 8 - typical loess, 9 - loess-like loam, 10 loam, 11 - sandy loam, 12 - sand, 13 - rock debris, 14 - pebble, 15 - laminations, 16 - angular unconformity, 17 - washout, 18 base of section, 19 - borehole, 20 - stony loess, 21 - tectonic jointing and cleavage in loess stratum, 22 - gliding planes, 23 - archaeological horizons, 24 - bone-bearing horizons, 25 - plant macrofossils, 26 - TL dates in ka by Shelkoplyas (KTL indicated index), 27 - normal polarity, 28 - reversed polarity, 29 - alternating polarity, 30- 31 - Matuyama- Brunhes transitional excursions and their local indexes. Sections with indication of height above sea level (m): 12 Fakhrabad, 9 - Kalai Melik, 9 - Guli-Kandos, 10 - Sa'ng-lak, 11 - Y a m 4 Site, I - Karatau, 13 - Kairubak, 14 - Chashmanigar, 15 Kbumtio (Kunitek). 16 Lakhuti. 17 - Karakchi.

-

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This Page Intentionally Left Blank

1 Is

Pleistocene loess, with the Upper Paks loess. The structure of the Upper Pleistocene portion of the loess sequence coincides in every detail with that of the Hungarian loess, including the thinnest humus interbeds, the Tapiosuli loess, in the Black Sea area. It should be mentioned that our correlation of climatic stratigraphic units of the loess zone with the oxygen isotope stages (Table 4.1) coincides with the earlier correlation of Kukla (1977).

4.2. Loess in Asia In this section we shall give a brief discussion of the loess blankets in Soviet central Asia and in northern China. Although its detailed study has been started only recently, progress has been fast. The introduction of palaeomagnetic and thermoluminescence techniques into practice has made it possible to dissect minutely extremely thick and practically “dumb” soil - loess sequences and to correlate them on a global scale. A lot of work has been carried out by Dodonov, Lazarenko, Penkov, Ranov and others in the southern Tadjik depression, whose sections have been taken as a standard for climatic stratigraphic units in Soviet central Asia (Fig. 4.4). Much progress has been achieved for the last ten years by Chinese researchers of the Pleistocene. Table 4.2 summarizes the data of the Soviet scientists derived from the monograph The USSR Stratigraphy. Quaternary System (Minina and Lazarenko, 1984), the reports of the International Symposium on the Neogene - Quaternary Boundary held in Dushanbe in 1977 (Nikiforova and Dodonov, 1977; International Symposium. . ., 1977) as well as the papers by Dodonov and Ranov (1984) and Lazarenko (1982). The Pleistocene loess units in Soviet central Asia have earlier been related mostly to three complexes of river terraces, namely the Vakhsh, Ilyak and Dushanbe loess in Tadjikistan and the Nany, Tashkent and Golodnaya steppe loess in Uzbekistan, which for the most part have not been described from palaeontological point of view. At present the main object of investigation is the interstream loess sequences, the thickness of which ranges from 15 to 200 m. The most representative of these are found in western foothills of the Pamirs outside the city of Kulyab and in the vicinity of Tashkent. Over 30 sections have been studied palaeomagnetically in every detail by Penkov, Semyenov, Toichiev, Yeroshkin and others. At the same time up to 9 - 10 excursions have been observed there in the Brunhes orthomagnethem. About 100 thermoluminescence dates of the sections and excursions have been obtained, out of which 60 have been made by Shelkoplyas et al. (we have used only his estimates in this part of the book). The degree of reliability of KTL dates is different. For instance, we have dates of 650 k 80 ka below the Brunhes - Matuyama boundary as well as 710 f 80 ka and 790 f 85 ka above this boundary (Fig. 4.4).

116

Table 4.2. Pleistocene climatostratigraphic units of Central Asia _________

7 -

Upper Malan loess

Upper loess. TL 17-47 rl excursion, TL 21 t 4

1

Paleolithic

5C

Soil III/IV. Obi-Garm

__

Travertine of the New Cave Th/U 68 - 75 and 98

Soil 1 TL 110

I

I

Soil IV/V, Lakhuti tools

5e

Breccia

Salawusu

Middle Paleolithic

Loess. r2 excursion. TL 120- 130

%I

Cave section of Choukoutlen ( = Zhoukoudien)

Northern China (Luochuan section, Shaanxi)

(Southern Tadjikistan)

Dingcun fauna

Loess, r3 excursion

7

( I ) Breccia Soil 11, TL 210 Sorex squornus

Soil V/VI

1

1

[

Loess

1

Soil VI. r k , TL 250

!

( 2 ) Travertine, Th/U 230 i 25

(3) Breccia, Th/U 256 i 50 Soil 111. TL 250

TL 290-310, ~ . t 300 . t 56 Loess 4

Loess, r4b. TL 280 - 300

Th/U 380- 420,Atluropodn,

Soil VII, tripartite(?) Excursions r5, TL 320

12

16,TL 380

16

+

* 35

8 - 9. Breccia. Cool steppe

Loess 5

50

(marker)

Loess 6

Loess Soil VIIl

I7

____

-1--

Loess and soil sequences

Forest - steppe

Soil VI

(12) Small breccia Dicerorhinos

Soil VII

R/N reversion (member 14) Longgushan formalion (members 14- 17) Gongwangling fauna: Ailuropoda, Sfegodon, Homo erecius lonrionensis

(13) Breccia-sand

Loess, B/M reversion

1

22

i

Erosional Dhase Lakhuti I1 fauna

Loess 8

I

Loess 9, erosional

_^--

Kuldara artephacts

Loess, n2 excursion Charh manigar Lakhuti I fauna

1 :I

I Loess I S , erosional phase

rErosional phase

117

The KTL date of 100 t- 15 ka for the second loess layer, counting down from the top, seems also t o be too high. On the whole, however, the dating of the sections seems to be quite representative, and the results of their correlation with the excursions are close to those obtained for the Black Sea area and the Ukrainian sections. In addition to palaeomagnetic and T L references, the correIation has also been based on the finds of Palaeolithic artifacts from certain soils (Dodonov and Ranov, 1984). In 1977 the author of this chapter of the book took part in research expeditions in Tadjikistan, where he had the opportunity of comparing the structure of the Tadjik loess sequences with these of the Sea of Azov region and the southern Ukraine. These proved to be very similar t o each other. Like in the Black Sea area, the palaeosols in Soviet central Asia form soil complexes consisting of two or three close or superimposed soils, the lower of which is brightly coloured. The soil complexes are separated from each other by loess horizons, whose thickness increases from bol tom to top. The features of palaeosols also change along the section. Down the section the soils (“shokh”) are “stony”, red in colour; they formed in the hot subtropical environment of broad-leaf forests and shrubs. The middle part of the complex is represented by meadow-steppe reddish brown soils. On top of these are greybrown dry steppe and semi-desert soils. On the whole, one can see through the sections a distinct rhythm of climatic fluctuations with repeated cycles of relatively humid and sharply arid conditions, which are associated with lower soil layers and loess, respectively. These features of the loess sequences in Soviet central Asia have also been noted by Popov et al. (1984). The reference level for correlating the Soviet central Asian loess deposits is the Brunhes - Matuyama boundary located under the roof of the tenth soil complex (Table 4.2), which evidently corresponds to the Upper Shirokino soil in the Ukraine. At the base of the tenth soil complex a regional angular unconformity has been found. The beginning of the erosional unconformity coincides with the bottom of the 10th and 12 - 13th soils, which, judging by the presence of the normal-polarity zone ( = Jaramillo?), corresponds to the Urzuf - Shirokino pedocomplex in the Ukraine. In conglomerates fixing the erosional unconformity a mammalian complex (Lakhuti 2), including Microtus sp. and Allophaiomys sp. of the Karai-Dubina type, has been discovered. The second indicator for correlating the loess sequences in Soviet central Asia is the fourth or the fifth soil complex after Lazarenko (1982) and Dodonov and Ranov (1984), respectively. (These authors suggest different numbers for the soils; therefore in Table 4.2 soils have been assigned two numbers, the first corresponding to Lazarenko’s nomenclature and the second, to Dodonov and Ranov’s.) There, in a number of sections the pre-Mousterian stone inventory ( = Lakhuti I) has been recorded, which, according to Dodonov and Ranov (1984), resembles the Micoquian culture. In the loess horizon that lies above the fourth (fifth) soil, a double zone of reversed polarity with KTL dates of 120 k 11 ka to 130 ? 25 ka has been observed. This excursion ( r 2a - b by the local nomenclature) most probably corresponds t o the Blake event. This allows us to correlate the fourth (fifth) soil with

I18

Priluki soil in the Ukraine and the underlying loess together with Karatau excursion (r3), with the Tyasmin loess. The lower interval of the Brunhes orthomagnethem is not so well correlated with the European sections. Excursion r5a in the roof of the seventh (eighth) soil is correlated with the Emperor excursion and excursion r5b at the lower boundary of the seventh (eighth) soil, with the Ureki excursion. The sixth soil of the Tadjik section corresponds to the Potyagailovka soil, and the eighth soil, to the Tsokur soil. Excursions r5 and r6 in the Lakhuti section seem to be an analogue of the Yakhno Event from the Don loess in the Black Sea area. The majority of the TL dates seems thus to be “rejuvenated”. The climatic stratigraphic Pleistocene units in northern China are described in the paper by Chinese scientists presented at the INQUA Congress in Moscow in 1982 (Liu Tung-sheng et al., 1982) and at the 27th International Geological Congress held in Moscow in 1984 (Huang Pei-hua, 1984, Liu Tung-sheng et al., 1984). Climatic chronological divisions of the Pleistocene in northern China can clearly be seen in the thoroughly studied section of the karst cave in Choukoutien near Peking (39”30‘N and 116”E), which includes all the Pleistocene deposits. The section in the Peking Man’s cave has been carefully described by Liu Ze-chun (1982). Here, the 47-metre thickness of deposits can be divided into 17 layers, of which the 4 lower layers in the narrow karst crevasse refer to the Matuyama orthomagnethem. They are identified as the Longgushan suite. The thirteen upper layers are included in the Brunhes orthomagnethem and form the Choukoutien suite. The Choukoutien cave was permanently inhabited by the ancient man, Homo erectus pekinensis. The section of the Peking Man cave is represented by alternating breccia beds containing huge limestone blocks and calcarious tuff, sand, ash and breccia beds of fine-grained texture. In the breccia of psephitic texture there have been discovered the pollen and faunistic remains of cold arid steppe and forest steppe environment with spruce and birch. The fine-grained breccia layers formed in the forest and forest - steppe environment as shown by pollen and faunal evidence. The forests included broad-leaved species such as Symplicos, which grows now only in southern China and the fauna was represented by relevant animals, including the southern Chinese Aifuropoda, Hystrix, Bubalus, orang-utan and so on. The geographical shift of natural zones, as found from thaphocenos and mineral and chemical composition of sediments, seems to be in the range of 2,000 km. The section in the cave is well dated by four independent methods, which makes possible its direct correlation with isotope stages (Liu Ze-chun, 1982). The loess sequence of northern China and first of all of the Ordos Plateau in Shan-si province is well known. The thickness of the Huantu formation (= loess s. str.) reaches 15 to 200 m and in some places even 300 m. Lithologically this sequence is divided into three suites: the Pliocene red Wucheng Loess, the Pleistocene brown Lishih Loess and the Late Pleistocene yellow Malan Loess. Each assemblage includes a great number of palaeosols, which are counted down from the top. The section at Louchuan in Shan-si province, which is 135 m deep, has been studied palaeomagnetically by two groups of scientists. According t o Heller and Liu Tungsheng (1982), the Brunhes - Matuyama boundary is located between the eighth and seventh soils, while the tenth to twelfth soils occur in the Jaramillo event. However, according to Zheng Hong-han (Quaternary Geology. . ., 1982, p. 59-65), the

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Brunhes orthomagnethem is thought to begin with the reversal at the level of the twelfth to fourteenth soils. Since other information for dating the Louchuan section is absent, the stratigraphic location of its soils and loess remains uncertain. It should be mentioned that the upper portion of the Nihewan suite with Equus sanmeniensis, which was thought to lie below the Wucheng Loess, seems to be located in the lower part of the Brunhes orthozone (see palaeomagnetic data of Li Hua-mei and Wang Jun-da, 1982). As can be seen, Table 4.2 suggests an interpretation of the Chinese information on the Lishih loess from two sources. According to Liu Tung-sheng, the Jaramillo Event is located within the Wucheng Loess, whereas the version of Cheng Hon-hang (Quaternary Geology and Environment of China, 1982) shows it within the lower Lishih loess. When comparing the data, we shall further rely on Liu Tung-sheng's version.2 Table 4.2 shows that the beginning of the conventional Pleistocene (below the Sicilian regional stage, 1.15 Ma) is marked by an erosional unconformity in all three sections. In Soviet central Asia it is represented by the Lakhuti conglomerates with palaeomagnetic age of 1.O to 0.8 Ma. In the Hwang H o valley this is an incised layer dividing the Wucheng and Lishih loesses and in the Yan-shan mountains this is an intense karst formation. The age of this phase (1.15-0.9 Ma) coincides with that in the Alpine, Carpathian and Crimean foothills, i.e. with the beginning of the Diluvium after Penk. This interface was also marked by drastic changes of successive mammalian faunas and the appearance of the genus Homo (Homo erectus), which is manifested by cultural stone material in Vallonnet, Azykh, Kuldara, Konvalin and Choukoutien XIII. From climatic aspects, the erosional unconformity developed during a period between two severely cold stages (the 15th and 9th horizons of the Lishih loess), which was very warm and humid (the Shirokino pedocomplex in the Ukraine, the Chashmanigar pedocomplex, the Choukoutien karst). These cold stages were evidently associated with the Xixabangma glacial stage recently established by the Chinese researchers in the Himalayas. According to Shie and Zheng (see Huang Pei-hua, 1984), on the slopes of Mount Xixabangma at a height of 8.012 m the glacial deposits are underlain with lucustrine deposits with fossil leaves of the oak (Quercus), which now grows 3,000 m below this boundary. The conclusion has been made that since the Xixabangma glacial stage, i.e. 3.0 to 0.9 Ma, the Himalayas have risen by 3,000 m. Similar information has been obtained by the Chinese scientists for the Tibet plateau (The Evolution of Eastern Asian Environments, 1983) and the Soviet researchers for the Pamirs (Pakhomov, 1982). The next stage was a complex thermochron at the end of the Matuyama orthomagnethem documented in Soviet central Asia by a transient mammalian fauna Lakhuti 2 with the first Microtus and the late Allophaiomys and in northern China, by the appearance of the southern Chinese fauna and flora (Huang Pei-hua, 1984). The upper portion of the Vakhsh loess in Tadjikistan and the sixth loess, which is thickest in the Lishih formation, evidently corresponds to the Nienxionla glacial in the Himalayas (Huang Pei-hua, 1984), which in turn is comparable with the Don glaciation in eastern Europe. According to Liu Tung-sheng et al. (1985), Fig. 4.5

' On the contrary, Dodonov and Ranov (1982) give preference for the alternative version

I20

shows that at that time the annual temperature in northern China was 9" below the present one and precitation decreased to 200 mm. The seventh Tadjikistan paleosol and the fourth and fifth Lishih paleosols are associated with two thermochrons at Choukoutien, which are clearly marked at the eastern exit of the Peking Man cave (layers 10- 8). The fission track, TL and Useries datings as well as geomagnetic excursions allow us to correlate this triple warming with isotope stages 15 - 11 (Huang Pei-hua, 1984). The seventh soil dated at 380 -t 50 ka in Tadjikistan, the fifth Louchuan soil, the Choukoutien travertines and soils Th/U dated at 462 k 45 ka formed within the warmest and most humid period during the Pleistocene (Fig. 4.4). The annual temperature in northern China was 5°C higher than the present one. According to Huang Pei-hua (1984), in northern Shan-si and the Beijing region the geographical environment was similar to the environment observed nowadays, 1,000 km southward, namely in the Yangtze

Annual average temperature

rC1

v -I===-

% s

200

Annual average precipitation 3w 4w 5w MXI 7w

E

I I

I

E l 1

U 3

m

GSC

I

Fig. 4.5. Upper part of Luochuan loess - soil sequence, Ordoss Plateau. Paleo-environmental parameters based largely on comparison with recent soils (after Liu rung-sheng et al., 1985). 1 - slightly weathered loess, 2 - medium weathered loess, 3 - strongly weathered loess, 4 - black loessial soil, 5 - calc-cinnamon soil, 6 - leached cinnamon soil, 7 - luvic cinnamon soil, 8 - brown cinnamon soil.

121

basin. This interglacial stage has been recorded in the Himalayas on the slope of Mount Xixabangma at a height of 4,900 m in lacustrine deposits with pollen of Magnolia and Carya. At present these species are found 1,500 to 2,000 m below this level. That means that the annual temperature in the Himalayas was then 3" to 6°C higher than nowadays and the mountains were 1,500 m lower (Huang Pei hua, 1984). Thus, between isotope stages 30 and 11 the average growth rate of the Himalayas was 3 mm per year and during the last 0.4 Ma, 4 mm per year. The Ilyak loess in Tadjikistan and the second, third, fourth and fifth horizons of Lishih loess seems to be associated with Jilong glacial stage in the Himalayas (Mount Qomolangma) and isotope stages 12, 10, 8 and 6. During the first phase of the glacial stage the arch of the Peking Man cave over the eastern exit floundered and the Peking Man moved t o the western exit. It should be mentioned that the cave at Choukoutien I had been a permanent habitat of the Peking Man for about 230 ka, and it bears many unique traces, revealing man's history, including the history of domestication of fire. This great event took place during a cold stage when the lower portion of layer 9 was forming, which corresponds to isotope stage 12 (Lui Tung-sheng et al., 1985). It should be noted that the thermochrons at the end of the Middle Pleistocene corresponding to isotope stage 9 and particularly to stage 7 are quite clearly pronounced both in Soviet central Asia and in China. During these periods the annual temperatures were 3" and 4°C higher than at present. However, the seventh thermochron was drier both in Asia and in Europe. It is also interesting to note that isotope stage 5 in the Soviet central Asian sections corresponds to three individual soil complexes (see Table 4.2). Isotope stage 3 is also associated with no less than three buried soils (inside the Malan formation), which are, however, pronouncedly weaker than the Lishih soils. Three details show that the loess zone was highly sensitive to changes in the global climate. According to Liu Tung-sheng et al. (1985), the transient belt between deserts and the steppe loess zone is particularly sensitive to such changes. Here, even a light decrease or increase in atmospheric precipitation causes the desert - steppe boundary to shift. The discussion of climatic events in the loess-desert zone leads us back to the acute problem about correlation between changes in atmospheric precipitation and temperature in the past. We recall that there are two contradictory points of view in regard t o this problem. Some authors (for instance, Markov et al. (1965) and many others in the past, Pakhomov (1982) and others at present) correlate the pluvial periods in Soviet central Asia with glacial epochs and the arid periods with interglacials. In particular, Obruchev's idea that the Asian loess formed in the warm environment was based on this point of view. Other scientists (Mamedov, 1982; Velichko, 1973; Dodonov and Ranov, 1984; Liu Tung-sheng et al., 1985) correlate loess deposits with glacial epochs and pluvial and soil-formation periods with interglacials. The information given in this and previous paragraphs leads us to the conclusion that the relationship between changes in temperature and precipitation in middle latitudes (35"N and 55"N) was rather complex and it is not explained by either of these two ideas. Let us now turn again to Table 3.4. Column 7 shows that the soil - loess sedimentary cycle consists of two stages and six phases. According to

122

Kukla (1977), the stages of the soil formation include four phases: (1) the forest phase; (2) the steppe phase; (3) the phase of dust storms, when key beds of eoline sands were deposited; and (4) the phase of pellet sands - downpours, when the soil accumulation on the slopes was retarded and the accumulation of the so-called redeposited soils was advanced (the latter soils are called “soil rocks” by Veklich and Sirenko (1976)). One wave of warming can be associated with three different soil layers, which form a soil complex. After Dylik (1957, 1969) and Washborn (1979/1980), the stage of loess accumulation also consists of three phases: (1) the humid solufluction phase; (2) the loess arid phase, and (3) the humid phase of washing-off. Thus, the temperature can change either to warming or to cooling, i.e. there are only two stages in the temperature cycles, the thermomer and the kryomer, whereas within the moisture course in time four stages can be identified within the climatic cycle: (1) the dry warm stage; (2) the humid stage; (3) the dry cold stage; and again (4) the humid stage. The humid stages within the cycle differ from each other by their degree of warmth. In order to verify these conclusions let us look at the data on fluctuations in the lake level in Central Asia. Investigations carried out by Devyatkin (1981) in the basin of the Great Lakes in western Mongolia (Lake Hirgis-Nur and others) have shown that these lakes underwent two pluvial transgressions within the climatic cycle. The first, great transgression developed when the climate changed from warm to cold, which is palynologically documented. The ages of the great transgressions of Lake Great Hirgis-Nur with KTL dates of 360 - 320 ka, 280 - 266 ka, 198 f 12 ka, 90 5 10 ka, 76 - 63 ka and 26 -t 3 ka coincide with the following boundaries of isotope stages 11/10, 9/8, 7/6c, 5e/5d, 5a/4 and 3/2. The second pluvials were minor and occurred at times when the climate changed from cold to warm, which is also confirmed by palynological evidence (Shilova, 1981), i.e. by an increase in the percentage content of pollen of dark coniferous species in the present steppe and semi-desert zone. Thus, within the climatic cycle the lake regressions also occurred twice, namely during the optimum of the interglacial and in the middle of the glacial; they were synchronous with the intense loess accumulation. These empirical conclusions allow us to suggest that within the climate cycle not less than four types of atmospheric circulation were observed: (1) the “normal” latitudinal zonal atmospheric circulation with an anticyclonic region above the desert and semi-desert zone; (2) the “hyper-cyclonic” circulation that forms under the conditions of warm ocean and gradually cooling continents; (3) the glacial anticyclonic circulation that appears with the advancement of ice sheets on the continents and the marine glaciation of the Arctic basin; and (4) the “hyper-monsoon” circulation that forms during late-postglacial time, when the ocean is still cold, while the continents are heated fast.

RCsumC (1) Over Eurasian territories covering a large area, the loess assemblage has gone through the same stages of development, which were determined by global climatic

123

changes. The thermoluminescence and palaeomagnetic reference marks show that loess horizons and interbedded soils formed synchronously in Europe, Central Asia and China. (2) The loess horizons were accumulated in dry and cold climatic environment. The buried soils of the loess assemblage over the plains were formed under more humid climatic conditions with temperatures close to the present ones. (3) The number of buried soils and loess horizons is generally greater than the number of isotope stages. Therefore, the “interglacial” stages of isotope scale often correspond to pedocomplexes incorporating several soils.

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Chapter 5 MIDDLE AND HIGH LATITUDES OF THE NORTHERN HEMISPHERE AS A MAJOR RECORD OF CONTINENTAL GLACIATIONS IN PLEISTOCENE TIME

5.1. Russian plain

Until recently the glacial assemblage sections have been used as stratotypes for climatostratigraphic division of the Pleistocene. Experience gained over the last century from the study of glacial assemblages of the Russian plain has been presented in the accounts by Yakovlev (1956), Markov and co-workers (1965), and Moskvitin (1967), who recognized four to seven glaciations. Till horizons related to marginal assemblages were used as the main markers for inter-regional correlation of sections. Pollen and spore analyses of intertill sequences show the existence in Pleistocene time of three or four interglacial epochs, namely, Likhvin, Odintsovo - Roslavl, Mikulino, and Mologa-Sheksna on the Russian plain. Some workers believe that pollen diagrams obtained for these epochs and, hence, climatic environments were fairly distinctive, and floras reflected, to a certain extent, some evolutionary steps (Katz, 1957, Grichuk, 1973, 1981, 1982, among others). However, there were contradictory viewpoints as well. They were clearly expressed by the distinguished Byelorussian scientist Voznyachuk (1965) who stated that of the three interglacials the oldest was the Roslav - Belovezha, synchronous with the Cromer s.str. of western Europe. Previously, the paleocarpologist Nikitin and geologist Dmitriev (1948) suggested an Early Pleistocene age for the glacial assemblage of the Don ice tongue. The findings of the Tiraspol rodent fauna with Mimomys intermedius in the Belovezha interglacial deposits near the village of Korchevo, Byelorussia (Voznyachuk et al., 1978) and later in the Roslavl interglacial stratotype near the town of Roslavl (Marginalformations. . ., 1985, p. 106, 183), favour the suggestion made by Voznyachuk. The works by Velichko and co-workers (1980) and Krasnenkov and co-workers (1983) furnished paleontological evidence for an older age of the Don glacial till as compared with the Dnieper till. The above has led to a complete misunderstanding about the stratigraphic position of glacial tills and interglacials over the European part of the USSR. The situation was further complicated by Moskvitin who in his latest works (1967, 1970, 1976) revised his earlier schemes. As a result, there appeared a lot of fairly complex compromise stratigraphic schemes where interglacial beds characterized by the same type of pollen diagrams belong to different stratigraphic horizons. Thus, in some schemes (Gursky et al., 1981; Goretsky, 1982; Zarrina and Krasnov, 1984; among others) the interglacial beds characterized by pollen diagrams of the Roslavl type were “divided” into the Shklov and Belovezha horizons, while the beds with diagrams of the Likhvin type were “divided” into the Bezhetsk and Likhvin horizons (Shik, 1981), and so on. The third meeting on the development of stratigraphic schemes for the Quater-

126

nary deposits of the European part of the USSR, held in 1984 and headed by 1.1. Krasnov, could not reconcile these disputed subjects. Evidently, stratigraphic subdivision of the glacial assemblage is frought with some objective difficulties. As stated above, the assemblage displays a mosaic structure, complex lithogenesis and lithofacies and is marked by glaciotectonic features and glacio-erosional phenomena. Goretsky (1970, 1980, 1982), Voznyachuk (1965, 1978, 1985), Kriger, Levkov, Spiridonov and others revealed a wide distribution of narrow erosiondenudation grooves formed during the advance and melting of ice sheets filled with thick strata of glacial drift, including interglacial and bedrock erratics. A number of known interglacial sections, including type sections (at Shklov, Korchevo), appear to be related to erratics. And, finally, it is quite evident that the paleomagnetic and radiometric approaches virtually cannot be applied to studies of the glacial assemblage. Consequently, the climatostratigraphic subdivision of the glacial assemblage appears to be extremely debatable. Table 5.1 is a synthesis of new climatostratigraphic data available for the East European plain. The table is based on such accounts as Geochronology of the USSR (Zubakov, 1974) and Stratigraphy of the USSR (Krasnov, 1984), as well as on the Proceedings of the third meeting on the development of stratigraphic schemes for Quaternary deposits of the European part of the USSR (Krasnov et al., 1986). Moreover, the stratigraphic schemes for Byelorussia and the Baltic Sea area were constructed with due regard for the data reported by Goretsky (1970, 1980, 1982), Vaitekunas and Gaigalas (1976), Yelovicheva (19791, Rylova (1980), Gursky and co-workers (198 l), Kondratiene (198 l ) , Makhnach and co-workers (1981), Voznyachuk (1965, 1985), Velichkevich (1982, 1984), Gaigalas and co-workers (1984), Yakubovskaya (1984) and others. In the stratigraphic schemes for the central and north-western Russian plain allowances were made for the data of Malakhovski and Markov (1969), Sudakova and co-workers (1974, 1977, 1981), Velichko and co-workers (1980), Arslanov and co-workers (1981), Shik (1981), Bolikhovskaya and Boyarskaya (1982), Goretsky (1983), Krasnenkov and co-workers (1984), Kozlov and Maudina (1985) and for the data presented in a number of such collected works as Complex studies. . . (1981), Moscow Ice Sheet. . . (1982), Puleogeography ofEurope . . . (Gerasimov and Velichko, 1982), Marginal formations . . . (1985), Problems of the Pleistocene (1985), Correlation of the Pleistocene. . . (1986). Interglacial correlation is based on stratigraphically related guide mammalian species (Voznyachuk, 1965, 1985; Markova, 1982; Krasnenkov et al., 1984), as well as on foetus and semen assemblages (Velichkevich, 1982, 1984). All in all in the Pleistocene glacial section there were established no less than twelve climatostratigraphic units. Some of them were further subdivided into minute units. The climatomers established were tied in to the oxygen-isotope scale using the projection of facies and stratigraphic changes of glacial drift into loess, and alluvial strata of the Dnieper and Don basins into marine deposits of the Black Sea area. The first appearance of periglacial plant assemblage which includes Selaginella seluginoides, Potamogeton filiformis, P. vaginatus, P. perforatus, and the like, is associated with lacustrine clays of the Vselyub formation in Byelorussia (Yakubovskaya, 1984), the upper Ivnyagi formation in the Moscow Region (Fursikova, 1984),

127

and the lower Daurnantai formation in Lithuania (Makhnach et al., 1981; Velichkevich, 1982; Gaigalas et al., 1984). This horizon is separated from true Pleistocene tills by the so-called Brest (Makhnach et al., 1981) or Vilnius (Kondratiene, 1981) interval with the "mosaic" pollen spectra that probably reflect a frequent alternation of forest and open grass and moss plant assemblages. Glacial erratics at the villages of Mikelevshchina and Korchevo, Byelorussia, Table 5 . I . Climatostrarigraphic units of rhe glacial part of Russian plain -

I

-

-

-1

1

5e

i I

Wikulino-Tiinoshkorii.hi TM, M T L 137

?

hl hl

10

? i

--I x

I b

9/11

11/13

,1

I

I

GrodnoTM

Aleksandriya TM Arv. moshuchmsrs

~___ L

I28

yielded t w o interglacial floral assemblages probably associated with the middle and late Brest stage. The assemblage from Mikelevshchina with Aracites johnstrupii and Brasenia bielorussica comprises up to 50% of Pliocene exotic species and gives evidence for a broad-leaved forest of the Balkan type (Yakubovskaya, 1984). The till of the first continental glaciation, named the Narev glaciation in Byelorussia (Voznyachuk, 1985), the Katler in Lithuania (Gaigalas et al., 1984), and the Kama glaciation in the cis-Urals area (Yakovlev, 1956), has been recently recognized in the Akulovo-type section near the village of Odintsovo, Moscow Region. The till horizon is the fourth from the top and named the Lika till by M.I. Maudina (Marginal formations. . ., 1985), p. 157). According to our correlation this glaciation is synchronous with isotopic stage 20 (or 22). Thus, as early as 800 - 900 ka BP the ice sheet spread south to 55” - 53”N. The Sivkovo periglacial flora is related to the initiation of the glaciation (Yakubovskaya, 1984). The Akulovo - Minsk - Paya interglacial, synchronous with the Taganrog thermochron, may have had two maxima. The interglacial is characterized by the presence of a broad-leaved forest containing such exotic species as Caulinia antiqua, Aldrovanda CJ dokturovski, Picea sect. Omorica, Pinus sect. Strobus, Celtis, and the like. Broad-leaved species spread up t o 60“N. Such representatives of the Mediterranean maquis as Hystrix came into Byelorussia from the Balkans (Voznyachuk et al., 1977). It is reliably stated now that the subsequent Byelorussian - Dzukiya - Don s.1. glaciation, synchronous with the Platovo kryochrone and isotopic stages 18 to 16, was maximal over most of the European part of the USSR. Till of its second stage is a single one near Lvov and in the southern part of the Don tongue. The till has been recently drilled in the Dnieper Valley as far as the town of Kanev. TL ages for the till in the Ukraine established by Shelkoplyas and co-workers (1986) fall in the range of 530-560 ka. In the Niemen basin the till has been dated by KTL at 560 - 600 k 67 ka, and reversed polarity zone has been ascertained in its upper part (Zubakov, 1974). In the Don basin the first, Lipetsk, stage of the glaciation is separated from the second maximal, Don, stage by the Moiseevo interstadial with cooler climate than at present (Krasnenkov et al., 1984). The “lower Oka” alluvium of the Chekalin section with a KTL age of 563 t- 70 ka may be also assigned to the interstadial (Fig. 5.1). Russet till with a KTL age around 230 - 280 ka in places directly overlies the Don grey till. The russet till is related to the Dnieper glaciation, hitherto considered as a maximal one (it has been found that the russet till transcends the Don till only in the Dnieper valley). Therefore the two tills have not been distinguished with the Don till only in the Dnieper valley). Therefore the two tills have not been distinguished within the Don and Dnieper tongues. Actually, they are separated by a lengthy time interval including up to four climatic optima corresponding to the four-partite Tsokur - Tambov - Gorod pedocomplex of the loess zone. This interval is still poorly known; previously it was assinged t o the “Great Likhvin”, whilst at drilling sites in the vicinity of the village of Odintsovo, the town of Roslavl, the village of Bibirevo it was referred to as the “Odintsovo interglacial” (Moskvitin, 1967, Shik, 1974). Correspondingly, the overlying - Dnieper - till was named the Moscow till here. Only now we can elucidate to a certain extent the true stratigraphic relations

I

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Fig. 5.1. Chekalin ( = Likhvin) key section of the East European Pleistocene. Compiled by the author from data of Faustov and co-workers (in Zubakov, 1973, p. 148); Sudakova et al. (1977); Vlasov et al. (1981); Bolikhovskaya and Boyarskaya (1982). 1 - sand, 2 - loam, 3 - clay, 4 - marl and gyttja, 5 - loess, 6 - buried soil, 7 - mammal remnants, 8 floral remnants, 9 - borehole, 10 - normal polarity, 11 - reversed polarity, 12 - arboreal pollen (AP), 13 - spores (SP), 14 - non-arboreal pollen (NAP). Sources of data on: lithology - Sudakova et al., polarity zones - Faustov et al., KTL - Shelkoplyas Vlasov et al.; pollen and estimated climate - Bolikhovand Ilyichev (in Sudakova et al., 1977); RTL skaya and Boyarskaya; OCT - author. ~

~

between the kryomers and thermomers of the central Russian plain (Table 5.1). The two-optimum Roslavl - Zhidin - Shklov interglacial is a stratigraphic equivalent of the Patraiy transgression of the Black Sea. Its lower, Glazovo - Belovezha - Muchkap optimum with the Tiraspol fauna corresponds to isotopic substage 15c and gives evidence of larch and broad-leaved forests dominated by some oak species Quercus petraea, Q. cerris and the like) accounding for 30% of the total, elm (Ulmus),up t o 22%, hazelnut (Corylus),up to 123%, and lime (Tylia tomentosa, T. cordata, P . platyphyllos), etc. The second - Pepelovo - Mogilev - optimum (substage 15a) is characterized by mixed coniferous and broad-leaved forests with hornbeam (Carpinus betulus), about 20070, and fir, about 15070, suggesting a more humid climate. The Pepelovo - Mogilev optimum is thought to be coeval with strata of the Smolensk

130

ford (near the village of Yakhny, on the Zapadnaya Dvina river), where Motuzko (Problems of the Pleistocene, 1985, p. 173) suggests the first appearance of the water-rat Arvicola mosbachensis. Velichkevich (1979, 1982) reports that the floras of the two optima were similar in composition with the presence of Cuuliniu sukaczevii, C. lithuanica, Brasenia borysthenicu, etc. The Podrudnyansky - Nizhninsky Rov cooling (substage 15b), separating the two optima, was a minor one judging from the flora, which is identical to the present mid-taiga forest. The distribution of such sub-Mediterranean species as Zelcovu, Ceitis, Tylia tornentosa, Picea sect. Omoricu, Quercus rerris, and the like (Moscow Ice Sheet , . ., 1982, p. 33) indicates that summer and annual temperatures during the Glazovo optimum were 3 - 4°C higher than at present, whilst in the Pepelovo - Mogilev optimum they were slightly higher than in Holocene time. The Belovezha and Likhvin interglacials are separated by the Oka-BerezinaDainava kryomer. The Bryankovo r-excursion and a KTL age of 483 f 59 ka (Sudakova et al., 1977) indicate that the kryomer may be coeval with the Malyikut kryomer and isotopic stage 14. The Oka ice sheet extended then to 54”N. Periglacial tundra steppe inhabited by the Late Tiraspol, Suvorov fauna with Dicrostonyx okaensis and Lemmus cf. simplicior (Aleksandrova, 1976) lay farther south. Lacustrine clay and gyttja of the Chekalin section at the Oka River yielding the Singil rodent fauna with Arvicolu mosbachensis-terrestris and Clethrionomys glariolus are the stratotype of the Likhvin interglacial s.str. The fauna and TL datings (KTL at 459 i 56 and RTL at 500 f 100 ka) allow correIation of laminated marl beds to the Middle Uzunlar and OCT 13. The Likhvin climatic optimum is represented by a fir and hornbeam forest with beech (Fugus), yew (Tuxus buccata), hemlock (Tsuga canadiensis), vine (Vitis) and Pterocaryu fraxinfolia. In bore holes the Likhvin beds are often taken for the Galich optimum and assigned to the “Odintsovo interglacial” (Moskvitin, 1967). The climate in Likhvin time was warm and humid. According to Grichuk and Kondratiene (1979), summer, winter, and annual temperatures were, respectively, 2 - 4”C, 1.5 - 3°C and 2 - 3°C higher than at present. Precipitation was 30- 40% higher (Fig. 5.1). The Kaluga ( = Kopys) kryomer probably corresponds to the Iksha ice sheet advance which reached the upper Volga area, 56” N. The subsequent (Chekalin, Grodno) warming is generally assigned to the end of the Likhvin interglacial, but Goretsky (1982) considers it as the first infra-Dnieper interstadial. In the upper Volga area it may be equivalent to the Kostroma interstadial, considered as the infra-Moscow (Moskvitin, 1967). All the beds contain taiga pollen spectra, broad-leaved species being common in Byelorussia and in Moscow Region. Goretsky (1980, 1982, 1983) convincingly proved that the Likhvin and Chekalin interglacials are synchronous with the Uzunlar transgression of the Black Sea. It has been found recently that the upper russet till developed in Byelorussia, in the Moscow Region, and in the upper Volga area, and named either the Moscow or the Sozh till (Moskvitin, 1967; Goretsky, 1980), or the Dnieper till (Voznyachuk, 1985; Breslav, Kozlov and Maudina (1985) and others, (see Problems of the Pleistocene”, 1985) may be related to a multistage (but single) glaciation. TL ages in the range of 230 - 280 ka and the reversed magnetic polarity of part of the intertill

131

loam (Odintsovo - Galich r-excursion with KTL at 280 f 30 ka) show that it could be correlative with isotopic stages 8, 9 and 10. This is consistent with the presence of the Volgian ( = Khazar) fauna including Dicrostonyx cf. simplicior and Lemrnus obensis in the pre-till beds (Sudakova et al., 1977). Paleontologically the Mikulino ( = Myarkin, Mga) thermochron recorded in tillunderlying clays with such guide fossils as Brasenia holsatica (Web.), Aldrovanda vesiculosa L. is best known. Its pollen diagram contains seven zones with the culminations, according to Grichuk (Gerasimov and Velichko, 1982, p. 92), of oak, elm and nut in zone 4, lime in zone 5, and hornbeam in zone 6. In Grichuk's reconstruction the boundaries of the broad-leaved forest zone during the Mikulino interglacial optimum are drawn respectively 5 - 5" and 1 - 2" farther north and south than the present one. Kondratiene (1979) reported that July, January and annual temperatures were respectively 3 - 5"C, 1 - 2.5"C, and 2 - 5°C higher than at present; precipitation was 30% higher. A natural zonation for the climatic optimum and quantitative climatic reconstructions were developed (Velichko, 1984). Unfortunately, the stratigraphy and age of the Mikulino interglacial remain uncertain. The stratotype, like reference sections of the Mikulino horizon in Byelorussia and in the Baltic Sea area, has not been TL dated.' Such data are available only for the Yaroslavl and Moscow Regions, and they are questionable. According to Sukachev et al. (1965) and Chebotareva and Makarycheva (1974), the best known sections of such Mikulino peat bogs as Cheremoshnik "B" and Dolgopolka are overlain here by the Kalinin till (Moskvitin, 1967). KTL ages for the till underlying the Dolgopolka and Cheremoshnik peat bogs are respectively 281 k 31 ka and 308 -t 32 ka. The overlying "Kalinin till" yielded values around 244 I27 ka and 209 f 22 ka (Sudakova and Ilyichev, 1974). KTL ages for the Mologa-Sheksna, according to Moskvitin (1967), gyttja of Lake Nero, presently correlated, according t o Sudakova and co-workers (1977), with the Mikulino interglacial, fall in the range of 114 - 95 ka. According to Moskvitin (1967), the type area of the Kalinin glaciation is the south-eastern part of Kalinin Region. During the Kalinin glaciation, the ice sheet reached the line connecting Rzhev, Dmitrov, Lake Nero, Kostroma. Geomorphologically, this zone of marginal formations certainly corresponds to the Warthe Glacial. The stratotype and the parastratotype of the Mologa-Sheksna interglacial are lacustrine deposits near the town of Rybinsk and those of Lake Tatishchevskoe near the town of Dmitrov, respectively. Recent integrated studies in the Lake Tatishchevskoe area show that (i) the KTL age of the Kalinin till is about 230 f 13 ka; (ii) the pollen diagram of the Mologa-Sheksna beds of the parastratotype (and stratotype) is similar to that of the Mikulino; and (iii) the climatic optimum of the Mologa-Sheksna ( = Tatishchevskoe) peak yields a somewhat older age about 110 ka and 100 f 10 ka. The above suggests that the Mikulino (in the sense of Moskvitin and Sukachev) thermochron and the Mologa-Sheksna ( = Nero, Tatishchevskoe) thermochron with its three optima can be correlated with isotopic stages 7 and 5 , respectively.

'

In late 1985 A.I. Shlyukov (personal communication) obtained a KTL age of 137 Mikulino beds in Timoshkovichi section, Byelorussia.

&

10 ka for the

132

However, a close similarity between the pollen diagrams of both Mologa-Sheksna and Cheremoshnik beds and that of Mikulino beds seems embarrassing. It should also be taken into account that, according to Grichuk (Moscow Ice Sheet . . ., 1982, p. 26), the Cheremoshnik “B” diagram resembles the Odintsovo, rather that Mikulino diagram. However, later Velichkevich (1985) discovered that a carpoid assemblage from the Cheremoshnik section, from the gyttja “B” inclusive, is undoubtedly younger than the Belovezha - Shklov flora. This may suggest that the Cheremoshnik “B” and Dolgopolka gyttja is related to some unknown thermomer with an age of 240 ka. Thus, now we have no reliable data to prove the age of the Mikulino interglacial, or to be more exact, “Riss - Wiirm” interglacials of the Russian plain. Table 5.1 assumes that the Mikulino ( = Timoshkovichi) thermochron may be a probable equivalent of the Mologa-Sheksna interglacial in the sense of Moskovitin (1967). The Valdai - Niemen kryochron represents a climatically complex interval with alternating coolings and warmings whose duration ranges from 1 - 2 to 5 - 6 ka (Zimenkov in Correlation of the Pleistocene . ., 1986, p. 133). The lower boundary of the interval remains uncertain. In compliance with 14C values, some authors (Arslanov et al., 1981; Zarrina and Krasnov, 1984) draw the boundary at the level of 70 ka. Other workers (Semenenko et al., 1981; Zimenkov, 1986), including the author, draw it at 95 - 110 ka in accordance with TL ages. In this case two upper Mologa-Sheksna optima become synchronous with the Chermanino and Tosno interstadials ( = Volozhin TM) in the sense of Zarrina and Krasnov (1984). The climate of the interstadials corresponds to that of middle and northern taiga with sporadic broad-leaved species. In many stratigraphic schemes, including that developed in Stratigraphy of the USSR (Zarrina and Krasnov, 1984), the Kalinin till is placed within the time interval of 57 - 50 ka under the name of Early Valdai till. However, as noted above, its young age has not been substantiated. Other workers (Auslender, Vigdorchik, Voznyachik, Kozlov, Zimenkov, among others) believe that the Kalinin till has been left by the last stage of the Dnieper ( = Moscow) Glaciation, while the maximal advance of the Valdai ice sheet took place between 22 and 18 ka. Ice reached the line Grodno, Vitebsk, Selizharovo, Vologda (The Latest Glaciation . . ., 1969). A periglacial zone was situated south of the line. The zone includes three or four interstadial warmings; the Grazhdansky Prospect - Krasnaya Gorka - Kyrvekyula warming with 14C ages ranging from 45 to 56 ka was the most important. Although the ice sheet might have been limited then by the Scandinavian mountains, the boundary of forest - tundra lay farther south, and the climate was cooler than at present (Zarrina and Krasnov, 1984).

.

5.2. Glaciated area in Western and Central Europe A comprehensive account of the climatostratigraphic division of Pleistocene glacial drift in western Europe and associated unsolved problems has been given by Kukla (1977) and Bowen (1978). This has considerably simplified our task, and the following is a brief discussion of the climatostratigraphic succession for western

I33

Europe. In compiling Table 5.2 for north-western Europe (GDR, FRG, Netherlands and, partly, for England) we used, in addition to the data of Kukla (1977) and Bowen (1978), the data of Van der Hammen et al. (1967, 1971), Erd (1970), Menke and Behre (1973), Zagwijn et al. (1971), Zagwijn and Doppert (1978) for the Lower Pleistocene, plus those of Frenzel (1967), Mitchell and co-workers (1973), Zagwijn (1973, 1975, 1985), Brunnacker (1979), Shotton (1983), Cepek and Erd (1982), Table 5.2. Clirnatostratigraphic units of north-western and middle Europe OCT

1

Lesrno - Pocnan KM, l4C 13 - 22 Magdalenian

.

Brandenburg KM Moershoofd - Denekamp TM

2 b

?

Zwierzyniec ( = Konin) TM. TL 71.6-75.9. Micoaue-Prondnikian

,”

-

-

*

Sartowice loess, Torun? till TL 6 7 - 71. Preszeletian

Kaszuby - Malbork? KM TL 91 - I01 5 13. r (Blake) excursion

5d

Mammalian ages

c

Grudziadz? ‘“C > 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
~

~

~

~

’ I n Siberia shallow-water river floodings form numerous river arm and temporary water b o d ~ r sover flood plains; they are called “sor”, cf. saara (Finnish) meaning small river or river arm.

148

Buried valleys have been established not only over lowlands but also in the surrounding mountains (in the Urals, Altai, Sayan and over the Middle Siberian High Plateau). The Ininsk formation represents Altai buried valley sediments, their thickness reaching 100 meters. The formation of ancient buried valleys is concurrent with erosion - tectonic phase manifested both in the mountainous areas and on the Kara Sea shelf. This leads us to the conclusion that mountain formation coincided with marine regression. The erosion phase according to paleomagnetic and TL datings of the Altai Bashkaus conglomerates, is 0.8 - 1 .O Ma old. This boundary has conventionally been assumed to be the Lower Pleistocene boundary in Siberia. Quite a different hydrographic system existed in West and Middle Siberia in the end of the Pliocene. Its scarce traces represent the lower levels of the peneplain in the vicinity of the Yenisey on the Middle Siberian plateau with absolute heights of 190-250 meters. This ancient alluvial plain is represented in the upper reaches of the Ob by Sagarlyk suite (Zudin et al., 1982) and by Razdol'e beds of Kochkovka formation (Adamenko, 1974). The Late Pliocene alluvial plain developed into VIIst fluvial terrace above the flood-plain - the Khudonogov terrace, which is a morphological feature of the present-day Yenisey (Fig. 5.6). Late Pliocene - Early Pleistocene alluvial plains at 56 - 60"N yield spore-pollen spectra containing up to 30% of exotic species like Pinus Sect. Strobus, Picea Sect. Omorica, Tsuga, Juglans and the like. This, together with red beds of Pliocene pebbles and the composition of Razdol'e fauna of small mammalian (near the Ob) of Villanyia, Allophajomys pliocaenicus, Lagurodon pannonicus, Pitymys ex gr. hintori-gregaloides, Pliomys kretzoi and others (Zazhigin, 1980), would indicate a warm climate with mild winters (Fig. 5.7). The first indications of cooling are recognized in the gray alluvial member of the Yenisey Khudonogovo terrace (56"N), yielding spore-pollen spectra containing Botrychium boreale, Betula nana, in the upper reaches of Ob, in clay near Tishinka village, near Barnaul, Kostitsina et al. (1966) and later Ponomareva (Zudin et al., 1982) found unique spore-pollen assemblages, having no equivalent. They indicate the coexistence (or frequent changes in time?) of plants which are now known as indicators of tundra environment, namely Betula nana, Lycopodiurn pungens, Selaginella sibirica, etc.) together with plants typical of forest and steppe zones like alder, elm, hazel, lime, Kochia and others. The Tishinka - Khudonogovo kryomer is taken as an equivalent of the Port-Katonian (Table 3.3). Kargat sands, yielding Eguus ex gr. siissenbornensis, according to Bukreyeva (1968) formed within a forest/steppe zone. They constitute the lower unit of the paleo buried valley section. Kargat flora in its composition is the same as plants living nowadays, while valley birch forests were abundant in broad-leaved trees like elm, hazel, lime. Obviously the climate was warmer and more humid than the present one. The age of basal strata of the Yenisey buried valley alluvium was determined by TL as 790 k 85 ka (Grechin, 1975). Complete reconstruction of natural zonality is believed to coincide in time with the formation of Ubinsk clays over the Barabin Lowland; the latter encompasses Berezhkovo beds in the upper reaches of the Yenisey and Yerestnaya clays in the vicinity of the Ob yielding Archichiskodon meridionalis tamanensis, Paracamelus alutensis and Equus sanmeniensis. This correlation, though, seems not so certain,

I49

since Yerestnaya beds contain Late Tamanian fauna, while Ubinsk clays yielded Bakinian ostracod fauna according to Kazmina (Zubakov, 1972a). Bukreyeva (1968), Kostitsina et al. (1966), Volkova et al. (1984), Zudin et al. (1982) showed that swampy-steppe landscapes with tundra species (Saxifraga caespitosa L., S. oppositofolia L., Rubus hamaemorus L., Selaginella sibirica and others), steppes and swamps (green and liverwort mosses) covered West Siberia along 55 - 56"N in Ubinsk - Sagarlyk time. If those landscapes are considered as tundra, then in Ubin time landscape zones shifted southwards by 9 - 10" of latitude and mean annual temperatures decreased by 7 - 8°C. The oldest of the seven tills in the glacial zone, still nameless,3 is equivalent to the Sagarlyk - Ubinsk kryochron. It was recognized by Cherepanov (personal communication) in a drillhole in the lower part of the section of an ancient buried valley in the mouth of the Irtysh (61"30"). The lower of the two Podkamennaya Tunguska tills in the

Fig. 5.7. Tentative reconstructions of the Pleistocene summer temperature in North West Siberia. :Cornpiled by the author from the data of Zubakov, 1972a, b; Kaplyanskaya and Tarnogradsky, 1974 1984; Volkova, 1977, 1984; Svitoch et al., 1979, 1981; Vasilchuk et al., 1984).

It was called Mansi by Arkhipov (1986).

150

section of the drill-hole IB near Lebed Settlement on the Yenisey (62”N) is likely to be of the same age. It was described by the author (Zubakov, 1972; figure 15). Lebed’ till should be a little younger than the TL dating, yielding 790 -t 85 ka, obtained by Il’ichev for basal pebble beds of the buried valley in the Podsopochnay River basin on the right shore of the Yenisey (Grechin, 1975). These tills by their stratigraphic location correspond t o the Kama glaciation, recognized by Jakovlev (1956) by a moraine in the old canyon near Solikamsk in the western Urals (59’30“). Thus, the edge of the Kama - Lebed Ice Sheet seems to roughly coincide with the boundaries of the maximum Samarian glaciation. There was Ubinsk Lake basin in front of the ice edge, the waters from the basin penetrated along the paleoOb deep valley as far as the foothills of the Altai, where they deposited “stony blue mud” of the “C” suite (by Pravoslavlev) which are now recognized as Yerestinaya beds. The lower boundary of the congelifluction-sor Fedosovo - Abalakovo formation formed concurrently over the interfluves. Judging by the whole volume of available data (TL analysis and faunistic evidence) the first Siberian glaciation can be correlated with the 20th or 22nd isotopic stage. Volkova et al. (1984) state that the Ubinsk kryomer is 1.8 - 1.6 Ma; that conclusion, however, seems rather groundless. The Talagayka - Teletskoye - Early Sergeevka thermomer corresponding to 630 k 27 - 800 ka by TL datings, and having Brunhes - Matuyama reversion (transition) layer at the top, yielded remains of Vyatkino mammalian assemblage with Equus srenonis cubbulus - Cervus eluphus and rodents Allophaiomys pliocaenicus, Pitimys hintoni, Mezosiphneus sp. (Zudin et al., 1982). Volkova et al. (1984) distinguish three climatic episodes for this period. Their estimated A T is 1.5 - 2°C for the first optimum, and in the remaining two the climate was harsher than at present. Kaplyanskaya and Tarnogradsky (1984) recognize the first appearance of the permafrost in the lower reaches of the Irtysh at the end of Talagayka time. The till of the second glaciation of the West Siberian Lowland was described by Epstein and the author by the core obtained in the drill-holes from the mouth of the Podkamennaya Tunguska in 1957 (Zubakov, 1972a). For the Ob it was described by Zakharov in 1961 by cores from drill-holes in the mouth of the Severnaya Sosva, near Shaytansky Cape. This moraine on the right shore of the Yenisey is TLdated as 500 f 60 ka (Grechin, 1975) while on the Belogor’ye hills on the Ob it is dated as 550 ? 100 ka (Arkhipov, 1984). Traces of this glaciation were found in the mountains as well, e.g. by Shchukina (1960) in the Altai and by Grosswald (1965) in the Sayan (see Zubakov, 1972). The Podkamennaya Tunguska- Shaytan glaciation correlates with lacustrine clay of Semeyka Suite of the Paleo-Irtysh valley, which was first described by Sukachev in 1910- 1934 (he called it “greyish loam”), which yielded in 1969 (Nikitin) the “greyish loam flora” of the “A” stage containing Seluginellu seluginoides Link, Nujus flexilis Willd., Betulu nunu L., Rununculus hyperboreus Rottb., Cochleariu urctica Schl. and others. Both investigators dated the “greyish loam” flora as corresponding to the Mindel Ice Age (Nikitin, 1965). This is confirmed by TL datings of the base of Semeyka Suite as 600 f 70 ka (Zubakov, 1972). As shown by the reconstructions made by Volkova et al. (1984) the forest-tundra zone shifted southwards down to 55”N in the

151

Semeyka kryochron, while the annual air temperatures were lower than the present ones by 7 - 8°C (Fig. 5.7). The next prolonged and extremely involved period in climatic history is documented in the sections of Tobol suite of the paleo-Irtysh valley, Laryak suite of the Paleo Ob, Kedrovskaya suite of the Paleo Tom’ and Turukhan suite of the Paleo-Yenisey valley. Most authors believe them to belong to the same interglacial, but the present author (Zubakov, 1972) recognizes at least three climatostratigraphic units there. The earlier thermomer (Vorogovo on the Yenisey, Chernyshevskoe on the Ob, Skorodum on the Irtysh) are described by Vyatka (Late Tiraspolian) mammalian fauna with Mimomys intennedius and Mammuthus trogontherii. The later thermomer (Panteleyev Jar on the Yenisey, Kalmanka on the Ob, TatarkaChembakchino on the Irtysh) contain Singilian fauna with Arvicofa kalmakiensis Zazh., Cervus ex gr. elaphus, Megaloceros, Palaeloxodon cf. antiquus (Zazhigin, 1980). Their TL-datings on I1 and I11 buried soils in congelifluction-sor assemblage near Belovo village on the Ob give their age as 536 -t 56 and 340 f 56 ka (Svitoch et al., 1978). KTL datings of Chembakchino alluvium in the Semeyka section (near Semeyka village) give its age as 380 +- 67 ka (Zubakov, 1972) and KTL datings on the Belogorsky-Materie section give 380 -t 90 ka (Arkhipov, 1984). Both interglacial members yield “seeded flora of diagonal sands” as it was called by Sukachev (1938). This flora, as shown by Nikitin (1940) and Nikitin (1965), contains such forms as Azolla interglacialica Nikit. ( = A . filiculoides), Bunias sukaczewii Kip. as well as abundant Potamogeton. It is interesting to note the presence of the mollusc Corbicula fluminalis there. The described interglacial alluvial members are divided by lenses of “greyish loam” which are of Sarchikha formation on the Yenisey, Krivosheino formation on the Ob, and of Demyanka formation on the Irtysh. Numerous tree stumps were found in situ by the author on the Yenisey, where in the sections of Panteleyev Yar (at 61”N) Sarchikha beds are recognized in the spore-pollen diagram by the peaks of Betula nana (almost 33%) and of Alnaster (about 26%), and the appearance of Lycopodium alpinun (L.) as well as a drastic decrease of fern spores (Zubakov, 1972a, p. 164). These lenses on the Ob and Irtysh are associated with seeded “Greyish loam” floras stage “B” mixed with arctic species. Krasnov found a horn of Alces latifrons in Sarchikha beds in the mouth of the Bakhta river, located well within the glaciated zone. Congelifluction-sor deposits in the Belovo section on the Ob correlated with Krivosheino beds were KTL - MTL dated as 410 k 40 ka. No tills corresponding to the Sarchikha - Krivosheino kryomer were found on the West Siberian lowland. Such a moraine, though, was found in the Altai Mountains. It was described by Shchukina (in 1960) who named it the Katun moraine and it was MTL-dated as 476 f 51 k a (Zubakov, 1974). This allows the Katun glaciation to be correlated with the 12th or 14th isotopic stage.4 Thus, the Vorogovo - Skorodumovo thermomer appears to be equivalent to the Roslavl thermomer, and the Panteleyev

The moraines dated as 413 k 52 ka recognized by Grechin (1975) at the Yenisey slope of Middle Siberian plateau are likely to belong to the same stage.

I52

Yar - Chembakchino t o the Likhvin interglacial. This is confirmed by isotope datings of animal bones (by 234U 238U) from Tobol sands of the buried valleys of the southern part of the eastern Urals as 400-450 ka (G.A. Shagalov, personal communication). The a foresaid makes it necessary to update the climatic curve of Volkova et al. (1984). Panteleyev Yar sections show evidence indicating that the Panteleyev interglacial, in turn, is divided into two thermomers by boundary strata of greyish loam with insitu tree stumps and a Bettula nana peak (Zubakov, 1972a, p. 162). There is a possibility that this Panteleyev silting stratum, like Sarchikha beds, is associated with a short-time continental glaciation of the Yenisey valley. This suggestion has been supported by the data of Borisov (1984) obtained in the Altai Mountains. Borisov found tills on the Kubadru river and in Bele, Teletskoye Lake, MTL dated as 304 f 40 ka and 340 k 40 ka. It thus appears that Tobol time corresponds to five climatomers of the Black Sea Region and to seven isotopic stages (from the 15th to the 10th). The maximum extension of the ice cap over the West Siberian lowland (down to 59"N) is believed to occur during the Samarov - Bakhta glaciation, with the Ural glaciers extending westwards and Taimyr/Putorana glaciers eastwards to 70 - 72"E. The northern part of the Middle Siberian plateau was also covered with ice. TL datings of the Ob till as 230 k 54 and 270 -t 56 and of the Yenisey moraines as from 240 f 26 to 312-413 k 52 ka allow them to be correlated with the 8th and 10th isotopic stages. In front of the ice cap there was the Yartsevo - Churym lake basin, whose contact with the ice is traced by a stretch of basin till. The lake stretched along river valleys as far as the Altai, where lacustrine deposits are TL-dated as 213 - 238 ka (Svitoch et al., 1978). The whole of the West Siberian Lowland in Samarovian time was covered by tundra or forest - tundra. Spruce taiga existed only in the Altai foothills. The fauna contained Volgan species (Khazarian) together with the representatives of the arctic and steppe fauna - Lemmus obensis, Dicrostonyx simplicior, Ovibos moschatus, Mammuthus primigenins pavlovae, Mammuthus trogontherii, Equus caballus, Lagurus lagurus et al. The landscape zone shifted southward by almost 1000 km while mean annual air temperatures were 9 - 10°C lower (see Fig. 5.7). The Shirta - Oplyvny Yar - Kartashevo thermochrone is a boundary, after which the present-day river network started to develop, though its corresponding level IV above-flood-plain terrace is morphologically manifest only in the foothills of the Sayan and Altai. Intermoraine members of the alluvium within the glaciated zone are TL-dated as 180-190 -t 40 ka on the Ob (Arkhipov et al. 1982) and as 246 -t 30 ka on the Yenisey (Grechin, 1975), while their corresponding IV buried Belovo soil in the steppes near the Ob is dated as 224 k 28 ka (Svitoch et al., 1978). Palinological evidence indicates the Shirta thermomer to have two climatic optima with the climate not harsher than the present one. The next Taz glaciation was not so extensive as the previous one. Glaciers reached Siberian ridges, its tongues extending as far south as 61'30" along the Ob and Yenisey. The congelifluction-sor assemblage of the low-level interfluve at the south of the lowland morphologically linked to the level IV over-flood-plain terrace of the Ob and Yenisey upper reaches is found to correlate with the glaciation. lt had two stages. The first stage till (the Yenisey stage) recognized by Zubakov (1972) was TL-

153

dated in the Yenisey basin (15 datings) as 200-240 ka (Grechin, 1975); its corresponding Khalapant Moraine in the Ob basin was not TL-dated (Arkhipov et al., 1982). Interstadial Strelnaya beds on the Yenisey (Zubakov, 1972) were TL-dated as 141 t 16 ka (Grechin, 1975). They correlate with Kormuzhikhanka beds in a well known Belogor'ye section on the Ob (demonstrated to the participants of the Moscow Session of the INQUA and IGC). Their base was dated as 130 f 3 1 ka and their top as 110 f 27 ka (Kaplyanskaya and Tarnogradsky, 1984). Evidence from pollen records indicates that Belogor'ye (62"N) at the time of the development of Kormuzhikhanka beds was located well in the sparse spruce forest zone and the climate there was colder than the present one. Arkhipov et al. (1982) overlooking the progressive age decrease of TL datings by about 10 - 20% incomparison with geomagnetic markers, mistook interstadial Kormuzhikhanka beds for interglacial Kazantsevo beds. This has resulted (as shown below) in an erroneous interpretation of the whole upper part of the section. During the second stage of the Taz glaciation, the Nizhnaya Tunguska stage, over the Middle Siberian plateau (Zubakov, 1972a) and Belogorye in the Ob basin (Kaplyanskaya and Tarnogradsky, 1984) glaciers advanced along the Yenisey and Ob valleys south to 63"N. It is still not clear whether the Siberian and Ural glaciers merged. Lake country and congelifluction-sor assemblage developed in front of the glacier which is reflected in the deposits of the sections along the Ob (Kiryas tributary and others). It is KTL-dated as 120 f 17 ka (Arkhipov et al., 1982). The till and fluvial glacial of the second stage of the Taz glaciation are dated on the Ob as 100- 110 t 17 ka (Arkhipov et al., 1982) and on the Yenisey as 120- 150 ka (Grechin, 1975). Taking into account the 10 - 20% age decrease of TL dates, these figures allow us to correlate the second stage of the Taz glaciation with isotopic substage 6a. It should be noted, though, that a number of researchers seem to share the same mistake considering the second stage of the Taz glaciation (Arkhipov et al., 1980) or its both stages (Troitsky, 1975) as traces of the Zyryanka glaciation. There is no good correspondence between the Taz glaciation and the Mayma glaciation over the Altai as shown by Suchukina, while Borisov (1984) claims that a correspondence exists between the Taz glaciation and the Chuya glaciation, TLdated as 145 k 13 ka. Its fluvial glacial streams moved to the foothill area where it is represented by pebble beds of the Biya Terrace and Bolshaya Rechka formation. These, together with their synchronous Kartashevo sands of the Irtysh river, yielded a mixed complex of mammals with ~ a m m u t ~ uprirnigenius, s Dicrostonyx simplicior and Lagurus lagurus described by Zazhigin (1980). The Mirnoe - Karymkary thermochron is the period when alluvium of the third above-flood-plain terraces of West Siberian rivers was formed. These terraces are well developed on the Ob and Yenisey, extending north from their upper reaches to about 63 - 64"N. The alluvium of this terrace and its synchroneous buried soils yielded early mammoth fauna and early Mousterian tools. The soils were TL-dated as 113 f 13 -91.7 k 11 ka (Svitoch et al., 1978), while the Altai alluvium on the Kuekhtanar river was dated as 90 - 109 ka (Borisov, 1984). The Yenisey section near Mirnoye village (62"N) and the Ob section near Karymkary village (62"N) are palinological type sections of the interglacial, the former revealing five stages in plant development with southern taiga species during the climatic optimum such as

154

the lime-tree and now extinct fern Osmunda cinnamomea L., together with an accessory of broad-leaved trees grass Stehria cf. holostea (Zubakov, 1972a). The latter yielded rich seeded flora Najas marina R. and Pofamogeton obfusifolius Mert. et Koch. (Nikitin, 1965). Fossil flora found in the Yenisey valley near Burmakino village (58”N) represents south taiga containing not only lime trees, but oak, elm and hazel as well. These and other data indicate a shift of landscape zones northward by 600-700 km (Zubakov, 1972). Volkova et al. (1984) estimated annual A T as 5 - 6”C, while the precipitation sum for the southern West Siberian lowland was estimated to be 600 mm. The Last Zyryanka glaciation over the West Siberian lowland is found to be a complicated climatic event. The extent and chronology of the glaciation are still debatable. Saks (1948) who was the first to discover this glaciation limited it to a narrow lowland stretch bounded by the Putorana plateau and the North Ural Mountains. Later some authors (Troitsky, 1975; Arkhipov et al., 1977, 1980) interpreted sub-latitudinal ridge topography between 65 and 70”N as a series of stadia1 marginal moraine belts, the age of the highest stage being estimated as 20-22 ka. Glaciers reached 62”N as glacial tongues along the Ob and Yenisey. Arkhipov and Astakhov (1980) claimed that the Kara Sea shelf and Yamal and Gydan peninsulas were the center of this glaciation. In front of the glacier there was vast Mansi or PozdnePreobrazhenskoye “Lake/Sea” fed by melt waters of the Lena jammed basin. The lake’s shore line was marked at 125 abs. m. The lake’s water was discharged through the Turgai Strait, Aral Lake, and Uzbey to the Caspian basin and further to the Black and Mediterranean Seas. This concept first proposed by Volkov et al. (1978) was further developed by Grosswald (1983), who then used it in his hypothesis of the Pan-Arctic ice sheet. Investigations carried out by the author seem to deny the foregoing. Troitsky (1975) and his followers ascribe moraines of the Taz - Yenisey glaciation TL-dated as 200- 260 ka to the maximum stage of the Zyryanka glaciation, or to its second stage (Nizhnyaya Tungusska - Belogor’ye) - TL-dated as 140 to 110 ka (Arkhipov et al., 1980). Their conclusion on the alleged Mansi Lake/Sea said to exist in the interval from 30-28 to 12- 10 ka and “great discharge” of melt waters through the Turgai Strait (Volkov et al., 1978) is based on a series of 14C datings of sections in above-flood-plain terraces of levels IV, 111 and 11. Some of these datings (not confirmed by TL dates) are erroneously considered younger. Valid datings, though, are not informative either; they do not prove the lacustrine nature of the basin. Most of these datings were obtained on congelifluction-sor features of Zyryanka Time developed in small recurrent lakes and ponds at various hypsometric levels. The water discharge through the Ubagan-Turgai strait seems rather doubtful and, finally, there is compelling evidence that there was no water discharge of Caspian waters through the Manych strait in the interval from 30 to 12 ka (see sections 3.2 and 3.3). Thus the whole Siberian spillway for melt waters seems to be a brilliant but speculative hypothesis. The analysis of materials on the Zyryanka glaciation discussed in the recent works of Arkhipov et al. (1977, 1980, 1982), Troitsky (1975), Kaplyanskaya and Tarnogradsky (1 984) has revealed an extremely complicated and twisted relationship of stages and interstadials with marine deposits, precluding their use as the basis for

155

dating of Zyryanka deposits. That is why the following discussion will cover only stratigraphic correlations of Zyryanka glaciation with continental features at the southern periphery of the Zyryanka glacian zone. Facial relationship of the alluvium of the Ob and Yenisey river terraces with the Zyryanka glaciation would lead to a conclusion on a three-stage structure of the latter (Table 5.3). The analysis of paleontological evidence suggests two mega-stages of the glaciation: the earlier Yermakovo stage and the later N'yapan stage; these are separated by Igarka megainterstadial. Sections in the lower reaches of the Yenisey, where the alluvium of level I1 above-flood-plain terrace (Igarka - Farkovo) joins the Yermakovo moraine N'yapan glacial horizon, reveals some facial changes with upper periglacial terrace beds (Konoshel'e beds), at some localities (lower than the lgarka river) it overlaps the terrace. That is why Igarka beds are considered to be the most reliable stratigraphic datum mark for the division of Zyryanka complex. Their age is 14C-dated as ranging from 32.5 to more than 50 ka (Zubakov, 1972a; 1974; Kind, 1974). Palinological records reveal a climate slightly cooler than the present one but in many aspects similar to it. The same characteristics of the section through the second level terrace of the Yenisey maintain also in the unglaciated zone (sections near villages of Pavlovshchina, Yukseyevo, Yermolayevo and others) and in the Angara Basin where the lower part of the terrace is recognized as Irkineyeva beds. The facial changes of the flood-plain terrace alluvial (or to be more accurate, of lower lacustrine - alluvial levels) with glacial complexes in the Ob valley have not been adequately studied yet and their descriptions by Lazukov ( 1 970 - 1972) and Arkhipov et al. (1977) are contradictory. The highest terrace in the present-day Ob lower reaches is the level Ill Karymkary terrace. Its upper part as seen in sections near the Bogdashkiny Hills is built up by congelifluction-sor deposits of the earlier stage of the Zyryanka glaciation with three hydromorphic soils. KTL datings of these beds referring them to 75 +- 10 ka ago (Arkhipov et al., 1982) allow the correlations to be established with isotope substages 5a, b, c, d and stage 4. The early Zyryanka congelifluction-sor assemblage is traced along the Ob, extending almost to Biysk, where TL dating indicates it to be 54.5 k 6 to 58.6 k 6 ka old (Svitoch et al., 1978). A similar age was obtained by TL dating for the moraine of the maximum Chibit stage of the Altai Mountain glaciation. The level I1 above-flood-plain terrace in the upper reaches of the Ob basin is well studied in Kargopolovo section by Pospelova and Gribidenko (1982), who recognized the Kargopolovo excursion in interstadial beds (Olbi - Laschamp) I4C-dated to be 42-43 ka. The same is true of the sections near Kolpashevo on the Ob and Khudyakovo and Lipovka on the Tobol (Zubakov, 1972a). The lower beds of the terrace (in all these sections) yield earlier types of Mammufhus fauna containing Bison priscus, Cervus cf. elaphus, Coelodonta antiguitatis and others together with spore-pollen assemblages typical of birch - spruce forests with numerous ferns. These beds were 14C-dated as ranging from 33 to 52 ka and older (Zubakov, 1972a). The upper part of the section is built up by a well-marked congelifluctionsor assemblage. Its lower unit is represented by varved silt at 57"N yielding sporepollen assemblages (zone 11), typical of forest-tundra. It is crowned by a boundary

156

horizon with series of greyish soils and with tree stumps in-situ (Larixsibirica Ldb., Picea obovata Ldb.) in the base of the unit I4C-dated as 30, 560 f 240 (LG-37)-30, 700 k 300 (GIN-126) years old. The section near the mouth of the Malaya Kheta River (69”30’N) at the lower Yenisey reaches displays Nyapan moraine (according to Troitsky and Kind, 1974) to rest on Karginsky beds, I4C-dated as 50 ka on the lower part of the section under the foot of the moraine to 37 ka. Thus the lower unit of the congelifluction-sor assemblage of the level I1 terraces of the Ob, Irtysh, Tobol was found to correlate with the N’yapan stage of the Zyryanka glaciation on the Yenisey 37 to 32 ka (Zubakov, 1972b), to Lokhpodgor stade on the Ob, also dated as 35 - 30 ka (Arkhipov et al., 1977, 1982) and to Akkem stade on the Altai (Borisov, 1984). The boundary horizon of Lipovka section displaying north-taiga spore-pollen spectra (stage 111) correspond to interstadial Novonazimovo warming some 30.7 - 26.3 ka BP. The climate of that interval was harsher than at present. The climate of West Siberia was the most severe and truly continental during the development of the upper part of the Late Zyryanka congelifluction - sor assemblage of the level I1 terraces 22 - 24 - 16 ka BP. This is confirmed by a large thickness of Elovka loess horizon (Zudin et al., 1982) and a pronounced adaptation of the animals of Late Mammuthus complex to almost snowless winters. The dimensions of the ice sheet over northern West Siberia is an extremely controversal problem lacking compelling evidence. Troitsky (1975) and Arkhipov et al. (1977, 1984) claim that it is in this time interval (22 - 16 ka) that the largest glacial advance over the lowland occurred. The authors together with other investigators (Saks, 1948; Lazukov, 1970- 1972) believe that the glacial advance was stopped by the foothills of the Putorana plateau and the Polar Urals. The ice retreated from the lowland approximately 16.5 - 15.0 ka ago during the Tab-Yakha - Uorankhalat interstadial when landscape zoning was very similar to the present zonation, while arboreal species on the Yamal Peninsula extended even further north than at present (Zubakov, 1972b). Sumin soils in the loess section correspond to this interstadial, displaying artifacts of the Late Paleolithic site Volchya Griva dated by I4C as 14.1 ka (Volkov et al., 1984). Later three short cooling episodes occurred, followed by two interstadials: Kokorovo, 12.94 - 12.26 ka and Taimyr, 11.77 - 11.25 ka (Zubakov, 1972b; 1974).

5.4. North-eastern Asia and Beringia

The region under consideration is a part of the so-called Beringia which in time of regressions served as a land “bridge” linking the two continents, representing an area of the most continental climate in the Northern Hemisphere, for which a hypothesis of glaciation metachronism has been proposed (Markov, 1938). Early data on the Pleistocene history of north-eastern Asia were reported by Kolosov (1947). The geological evolution of Beringia was a topic of two international symposia - Bering Land Bridge (Hopkins, ed., 1967); and Paieoecology of Beringia (Hopkins, ed., 1982). New information was published in the volume of

157

papers presented at the meeting on the development of stratigraphic schemes for the North East USSR (Quaternary deposits of the east o f t h e USSR - Biske, ed., 1982) and summarized in Stratigraphy of the USSR (Zamoruev and Petrov, 1984). Nevertheless, the climatostratigraphy of the Pleistocene section of the north-east USSR is still not clearly understood; the same is true of adjacent Alaska, since the geological history of both areas are very similar. In view of the heterogeneity in morphology of the region, a succession of the Pleistocene climatic events is considered separately here for three columns representing the Kolyma Highland, terrace-like features of the Bering and Chuckchee Seas, and Yana - Kolyma lake - alluvial plain (Table 5.4). Table 5.4. Pleistocene climatostratigraphic units of north-eastern Asia and western Alaska OCT

1

Mountains of North-Eastern Asia Cordeev - Sakhyn'ya KM, RTL 17-24 0

'2

-

-5

Lower Lena (?) T M l4C 26-30?

M

vl 2

1

I s k a t m - Napiowne - Donnelly KM Regression . . . - 100 m

Ulakhan-Kuel? KM

\

Mus-Khaya KM Tundra - steppe, ' j C I 2 - 22 Kurenakh-Salan T M

U s t ' Lora Amguema? Penrhina Bootthleger Cove? Canyon-Creek TM, ~

~

Nemkin-Tengelyakh T M RTL 37-44. Forest-tundra

Jana - Knlyma lowlands

Marine shore

~

Eva-Creek Ash, "C 2 56.9.1 F.1. < 80, Th/U < 80

Forest - tundra. "C 2 56 ~

Upper Vechernino- KukhtulZhigansk? KM R T L 5 0 + 13-60 16

*

- Kn,k

Valkarai Delta? KM

Malyk-Sien - Nomenkur KM RTL 135-148 i 40

G ' k a i l e n - Konnergmo'! Attarman Pelukian TM, Vc faram-zone (Buccelu rro,isk.vi). forest fauna. I4C > 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

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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

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1

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

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8

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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

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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.

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SECTION 111

THE HISTORY OF CLIMATE THROUGH THE PLEISTOCENE

This Page Intentionally Left Blank

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

--

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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

5a-c t

i12g

5d

Tyrrhen ( = Srrombus) Karangat

%

5e

6

.

.?

$5

WI

Small R - W lino)

(=

Miku.

‘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).

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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.

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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

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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.

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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?)

<

Q

m

-"9

TTL 56

Lr. 'ngrad

(

'4C43 I - 251

Kurgaloro KM TTL izn z lo?

Mirogoshcha

KM. 30.7

7

/I

Plum Point Farmdale T M ~

?

- 32.27 Cherr, tree - Tenaia K \ 1

Hengelo TM

KM (Ruigrkluft?) 38.7 - 4 0 5 KhornuT-Yuryakh -' Terekhtyakh - Berezavka

Maershoafd Upton Warren Nygaard T M

~

CB

-

14c 44.5

-

45 - 55

( = Stillfried

> 53 3

R rail?)

F.1.

<

80

nach -

ujyagos

KM ( = Ermakobo? Kharhgori?)

Denekamp - Kesseli- tiota-Alv T M ca 2 2 - 2 8 - 3 4

I:

Denezhkino ( = Errnako*o?)

Bogdashkin) Gor) T M . TL 75 t 10

Lake Erie T M

-

37.0- 3 8 . 7

1

Toino - Tarasovo Volozhin Karukula? Jonyeni, TM IJC 2 52

2

sradr

KM

:onsin

Part BruucePindale KM

Siegernde KM

Corha TM

Shenckoe

Late Pleniglacial

Undulyung T M 15.0-15 85

T M 15.2- 16.5

-

1

Nirningda

Putorana KM

Ostashkov KM

North America

Western Europe

North-Eastern Asla

-

^.

."

Lnigansi!

Earl) Pleniglacial -

1 I ?

N L A w

/I 5.0

8"O

PDB

4.5

4.0

VI

VIII

VII

3.5

IX

100 rl-

r27-3;;

7-3;

r41

' I

3

2

Fig. 8.7. Correlation of the Wiirm climatic events. 1 - oxygen-isotope stages and substages; 11 - oxygen-isotope curve in core V 19-30, Pacific Ocean (Shackleton et al., 1983); 111 - sapropels (SI -5) and ash layers - B , x , - ~ and , w,)in deep-sea cores from eastern Mediterranean (Bleschmidt et al., 1982; Cita et al., 1977, 1981; Thunell and Williams, 1983; Muerdter and Kennett, 1984); IV - sea level changes and coastal lines (Aharon, 1984, 1985; Harmon et al., 1983); V geomagnetic excursions (see Table 2.5); VI - loess- soil sequences in the Ukraine (Veklich, 1982), Czechoslovakia (Kukla, 1977), and Hungary (Pecsi. 1982; Marusczak et al., 1983); VI1 - pollen curve of Grande Pile peat (after Woillard and Mook, 1982), pollen zones designated by numbers 1 to 21; VIIl - two dynamic types of marginal parts of Wiirm ice sheets in the Northern Hemisphere: (a) - Arctic and West Siberian type and (b) - Huron - Europe type (after Zubakov, 1966). Till and intertill bed symbols (in VIII): E - D - Eem - Don, S P - Saint Pierre, Br I - I1 - Bradville tills, L - Leningrad till, P - T - Port Talbot I - 11, D - Danwich till, C - T - Cherritree till, P - P - Plum Point, Sk - Ka - time interval of the Skaerumhede and Karginsky marine suites; IX - maximum of Greenland ice sheet accumulation (Dansgaard et al., 1984). ~

-

237

nean (Vergnaud-Grazzini et al., 1977; Blenchschmidt et al., 1982; Cita et al., 1981). In the USSR its standard is the sequences of the second fluvial terrace above the flood-plain (together with top cover) in the Dniester and Don basins near the Palaeolithic camps in Molodovo (Inanova, 1977, 1985) and Kostyenki (Praslov and Rogachev, 1982).

8.3.2. The Early Wiirm, or kryo-orthoclimathem Sd - 4, 117- 62 ka In the Grande Pile section, the Early Wurm embraces pollen zones from 3 to 12. Zones 3 , 7 and 12 characterized by a low content of pollen of arboreous species (Fig. 8.1) are correlated with isotope substages 5d, 5b and stage 4, whereas zones 6 and 8, according to Menke and Tynni (1984), correspond to the Brorup and Odderade interstadials. Thus the double thermomer of Saint Germain is an equivalent of the double Vitachev soil in the Ukraine (TL dated at 57 to 85 ka) and the double soil of Saint Romain in Normandy (TL dated at 83 to 88 ka by Wintle et al., 1984). Investigations on the Bermudas tectonically stable platform have shown that the sea level was not higher than 15 and 20 metres even within warm isotope substages 5c and 5a (Fig. 8.2). On the slowly rising coast of the Mediterranean, substages 5c and 5a correspond t o the third Strombus terrace ( = Epityrrehen - Surozh) with two levels; y , dated by Butzer at 110 f 5 ka and y, dated at 80 k 5 ka (Bowen, 1978). The water exchange between the Mediterranean and the Black Sea within substages 5c and 5a can be confirmed by the presence of thin sapropels S, and S, in the cores from the Mediterranean with an age of 99 to 101 ka and 79 to 82 ka (Muerdter and Kennett, 1984). Thus, oxygen-isotope data and the data on the sea level show that the volume of continental ice was greater than at present even during the Early Wurm thermochrons. Within the glacial stages of the Early Wurm, i.e. substages 5d, 5b and stage 4, the sea level decreased by more than 20 m (Harmon et al., 1983). This fact shows that the formation of ice sheets proceeded in Scandinavia, Canada and Putorana. It is supposed that these stages correspond t o the tills of Malborg and Torun in the Vistula estuary (Mojsky, 1982, 1985) and the Yermakovo and Denezhkino tills in the Yenisey basin (Table 8.3). It should be mentioned that there gletcher ice has been found, which has been preserved straight from the Denezhkino up to our days buried under till and peat dated at 43,600 k 1,000 14C years (GIN 1896), the ice itself being of 50 ka 14C age (GIN 1892) (Quaternary Glaciation of the Middle Siberian, 1896, p. 34). In North America the edge of the continental ice sheet advanced during stage 4 and probably substage 5d as far as the Great Lakes, which is indicated by the presence of two Bradtwill tills separated by aleurites (Dreimanis and Goldthwait, 1973). However, the glacial advances southward during the Early Wurm Wisconsin were not the greatest. But to the north, within the Arctic and possibly sub-Arctic, the glaciation reached its maximum development just during the Early Wurm (Zubakov, 1972b, 1974; Boulton, 1979; Andrews et al., 1984; Vincent, 1984). The causes of this time lag of the Wurm glaciers can be understood considering the data on the Norwegian Sea and Greenland. The isotope analysis of the ice cores from borehole Dye 3 in southern Greenland has shown that the Greenland ice sheet

238 Core V30-97 Estimated summerlwlnter SST (OC)

6'80 (per mil) V30- 10 1K

55 1

5.0

45

40

35

30 1

C a C 0 3 (%) 50

70

60 h

A

A

,

80 .

.

A

A

Fig. 8.8. Downcore record of 6"O, the per thousand enrichment of " 0 , ( ice volume), estimated summer and winter sea-surface temperature, and CaCO, in northern subtropical gyre core V 30-97. Isotopic values are based on Uvigerina peregrina (circles) and three species referenced to PDB and corrected to Uvigerina: squares are Cibicides wue//ersfor$ ( + 0.64). diamonds are Pyrgo spp. ( - 0.51), and triangles are Melonis spp. ( + 0.40). Values from core V 30- 101 K (open symbols) were spliced with values from core V 30- 97 (filled symbols) by I4C control. Temperature estimates are based on transfer function derived by Ruddiman and Glover. The CaCO, analyses are precise to i 0 . 5 % . Isotopic stages 1 t o 8 are indicated in the column on the left (after Ruddiman and McIntyre, 1981b).

239

formed in the main at the end of the Riss - Wurm and during the Early Wurm, i.e. about 120-60 ka BP (Dansgaard et al., 1982), when the North Atlantic was still warm. According to Ruddiman and McIntyre (1981) and Streeter et al. (1982), the temperature of the bottom water of the Norwegian basin within substages 5d - 5b was 2-4°C higher than at present and the surface water temperature was the greatest within substages 5c/5b, i.e. about 100 - 90 ka BP (Fig. 8.8). This figure also shows that within substage 5a the winter and summer surface water temperature in the Norwegian basin was 5°C below the present one. Thus, following the Ruddiman - Mclntyre model and Dansgaard's data it can be stated that the ice sheets in high latitudes of the Northern Hemisphere were forming under climatically contrasting environment, when the oceans were still warm, while the continents were already cold, i.e. within substages 5d and 4 (about 110-60 ka BP). A considerable decrease in temperature of the Norwegian surface waters within both of the Early Wurm interstadials is evidently associated with a decrease in water salinity because of the melt water inflow and the development of persistent pack ice. An interesting phenomenon of the Early Wurm is the Pleistocene highest levels of the Early Wurm transgressions in the lacustrine basins between 35 and 50" N Lat. Thus, the maximum Early Khvalyn increase in the Caspian Sea level is dated by TL technique and the presence of the Blake excursion at I13 - 70 ka BP (Table 8.3), i.e. it occurred within stages 5d - 4. In Lake Sierls the thickest bed of lacustrine bottom mud from borehole LDW-6 is dated by the U-series technique at 122.3 t 3.6 ka t o 58.4 k 1.5 ka (Bischoff et al., 1985). In high latitudes of the eastern sector of the Northern Hemisphere, the Early Wurm is marked by the appearance of the so-called Oiyagos Edoma formation, i.e. by the accumulation of ground ice and the appearance of kryo-xerophytic Arctic steppes with their highly specialized fauna of the late mammoth assemblages, which adapted themselves to severe winters with little snow. All these facts permit the conclusion that the Early Wurmian climate was characterized by a strong atmospheric circulation (compared with that of other glacial periods), an increase in temperature differences between the ocean and land and an enhanced atmospheric precipitation in the Atlantic sector in both high and middle latitudes accompanied with more pronounced climatic continentality in the eastern Arctic and the sub-Arctic. It may be notice that in the Early Wurm volcanic activity and the geomagnetic field became more intense. Thus, substage 5d coincides with the double Blake reversed polarity event K/Ar dated at 113 ka BP. Thick ash beds of similar age have been found in the eastern Mediterranean. These are ash X5 - X6 from the Campanian volcanoes (Cita et al., 1981), ash in the sixth Hungarian young loess sequences (PeEsi, 1982) and in the Alaska loess sequences, where the Old Crow ash occurs above the Blake R-zone (1 13 ka) and is dated by the fission-track method at 120 ka (Westgate, 1982). During the time of the maximum ice growth of the Greenland sheet, i.e. 7 5 k a BP (Dansgaard et al., 1982), the greatest Late Pleistocene volcanic eruption of Toba on Sumatra dated by K/Ar at 73.5 t- 3 ka BP (Ninkovich et al., 1978) took place.2 The ash of this age has been found within isotope stage 4 in the

* Nishirnura has dated the tuff by fission-track at

100

+_

20 ka (Yokoyama, 1984)

240

Mediterranean cores tephra x2 (Cita et al., 1981) and in cores from the Indian Ocean. Such a coincidence of the phases of intense atmospheric curculation, ice growth, pronounced volcanic and geomagnetic activity shows that there was probably a causative relationship between them. It is possible that simultaneous intense in activity of all geospheres of the Earth was directed by an external factor, namely by changes in the Earth's orbital parameters.

8.3.3. The Middle Wurm - thermochron 3c, 62 - 42 ka The climatostratigraphic standard of the intra-Wiirm warming in the deep-sea sections is core KET 8004 (Fig. 8.9) in the Tyrrhenian Sea (Paterne et al., 1986), whereas in the continental sections it is the Grande Pile sections (Woillard and Mook, 1981) in the Vosges. The oxygen-isotope curve for KET 8004 allows one to identify within stage 3 eight light peaks, whose age has been determined from the sedimentation rate and correlation with five ash interbeds and turned out to be 60.5, 57, 52.4, 49, 44.4, 37.5, 33.4 and 3 1 ka (Patterne et al., 1986; Valladas et al., 1986). The same number of undated peaks has been recognized in the isotope curve for V19 - 30 in the Pacific Ocean. In the loess and glacial sequences of the Northern Hemisphere there are up t o five warmings forming two or three interstadials (Fig. 8.1). Some of the schemes give the upper boundary of the interpleniglacial as coinciding with the top of isotope stage 3 (Zubakov, 1972; Wysoczanski-Minkowicz, 1980), others present it as high as 22 - 25 ka (Van Hammen et al., 1971; Dreimanis and Raukas, 1975; Kind, 1974; Zarrina and Krasnov, 1984). KET 8004

P 1

2

3

4

0

20 1

2

-----

40

60 3

JIH

4

G

5

Isotopic -toges Iou~tier layers

Fig. 8.9. Comparison of 6 ' * 0 variations measured in core KET 8004, Tyrrhenian Sea, and dated Moustierian layers in Dordoyne sites, France (after Valladas et al., 1986). Ages of oxygen-isotopic events (light 6"O peaks) were calculated assuming a constant rate of sedimentation of reliably dated ash layers (Paterne et al., 1986).

24 1

The climatic optimum of the intra-Wurm warming was very short; it lasted only a few thousand years. According to I4C dates, its maximum took place 43 ka ago (Coope, 197) or from other sources 48 - 45 ka (Arslanov et al., 1980), 44,000 years, (Grn 6807) and 47 +- 1 ka (Laukhin et al. and Ivanova in Geochronology of the USSR . . ., 1980). On the whole these estimates have coincided with those obtained for deep-sea sections. As has already been mentioned, it is highly difficult to reconstruct climatic conditions for thermochron 3c (the Moershooft - Igarka therrnochron), because it includes positively interglacial deposits, which are known to be more ancient, but which, for some reason or other, have yielded younger I4C dates. Therefore, we cannot say how reliable are the numerous palaeoclimatic reconstructions for intraWurm time. It can only be stated that all available information shows that the climate in Europe and North America was unstable and cool, more continental and dry as compared with that of the present (Arslanov et al., 1980, 1983; Karrow and Warner, 1984; Dreimanis and Goldthwait, 1973; Zagwijn, 1974; Woillard and Mook, 1982). For example, a detailed spore and pollen analysis of peat soil in the valley of the Dziguta river on the Caucasian Black Sea coast made by Gey has shown that between 47 and 38 ka ago the climate was cooler than at present (Geochronology of the USSR . . ., 1980, p. 137). The results of the analysis of Coleopteron assemblages from 20 western European sections have led Coope (1975) to the conclusion that only during a very brief time interval of about 43 ka (evidently 49 2 2 ka BP) was the summer temperature higher than at present and the winter temperature lower. At about the same time of 44.5 +- 1.870 ka ago (LU 1050), climatic conditions of eastern Arctic “were less favourable than at our time” (Arslanov, Vereshchagin et al. in Geochronology of the USSR . . ., 1980). They came to this conclusion while thoroughly examining the habitat of the Terekhtyakh mammoth, whose frozen body and the accompanied rich Coleoptera assemblage (3 1 species) have been discovered in the lower reaches of the lndigirka river. Thus, the low sea level, which was 40 metres below the present one, as well as the data on soils of the loess sequences and the reliably dated pollen diagrams unanimously show that all over the Northern Hemisphere, including the Arctic, the climate of the interpleniglacial was cooler and more continental than at present. The climatic picture of the intra-Wurm interglacial (Karginski - Ponoi) turned out to be wrong because of a stratigraphic error.

8.3.4. The Late Wurm, 42- I3 ka The last kryorner, i.e. the time interval that witnessed no significant warming, can be considered t o begin some 40 ka ago. By that time the Laurentide ice sheet once more had filled the depression of Hudson Bay and entered the Great Lakes basin, its traces being the till of the “Cherry-tree” stage. The I4C age of this boundary obtained by analyzing wood from the Waterloo section, Ontario, is 40,080 & 1,200 years (Karrow and Warner, 1984). In Scandinavia it corresponds to the Haagesund stage. In north-eastern Asia this boundary divides the time of the existence of the Shandrin and Kirgilyakh mammouths, whose frozen bodies have been dated quite reliably and whose habitats have been examined to the utmost degree. The mam-

242

mouth found on the Shandrin river, a tributary of the Indigirska, I4C dated at 41,750 ? 1,290 years (LU 505) by its muscular tisues (Arslanov, Vereshchagin et al. in: Geochronology o f t h e USSR . . ., 1980) lived at the time of the intra-Wurm climatic optimum, when the boundary of the light forest reconstructed from the contents of the mammouth’s stomach and pollen diagrams ran 200 km to the north of its present position. A baby mammouth called Dima from the Kirgilyakh river discovered by Vereshchagin died 39,570 f 870 14C years ago (LU 718A), when the climate was colder than at present. This conclusion has also been corroborated by the analysis of the stomach contents of the horse’s body discovered in a mine gallery on the Selirikan, a tributary of the Indigirka, which was dated by I4C method at 38,590 _t 1,120 years (LU 506). The new data from core KET 4008 of the Tyrrhenian Sea are also indicative of a change of the warming for the cooling that occurred 42 (or 47?) ka ago. It is significant that this boundary can be clearly seen in the evolution of the organic world. According t o TL dates, in the sections of Combe-Grenal and Le Moustier, Dordogne a swift and definitive transfer from the Mousterian to the Late Palaeolithic occurred 42 - 40 ka ago (Mellars, 1986; Valladas et al., 1986). Before that time, in the Early Wurm, the Mousterian culture coexisted with the archaic Late Palaeolithic (the Szeletian ?, Egerian, the Volkov site in Kostyenki). The boundary of 42-40 ka can also be seen in the replacement of the forest rodent fauna with Apodemus by steppe assemblages with Citellus (Chaline In: Rozycki, 1982). The oxygen-isotope curves for ground ice drawn for both syngenetic ice wedges and organic-mineral rocks with polygonal ice veins of the Edoma complex proved to be extremely valuable in climatic reconstructions of the Late Wurm. This information embracing a time interval from 40 ka ago to the present time have recently been published by Vasilchuk et al. (1984, 1985); see Fig. 8.10. Isotope data (6l80 . . .-25.2%,) approaching these of the present ice veins (6l80 . . .24.6-27%,) have been obtained only for the samples of age 37-36 ka. Later within the time interval from 36 to 9 ka ago they were below the present values also

Fig. 8.1O.Oxygen-isotope curve along organic sedimentary polygonal ice wedges through the Zeleny Mys section, Kolyma river valley (after Vasilchuk et al., 1985). 1 - loam, 2 - vegetation peat remnants, 3 - reformed ice wedges, 4 - 7 - points of sampling for: 4 - radiocarbon analysis, 5 - oxygen-isotope analysis, 6 - pollen analysis, 7 - hydrochemical analysis, 8 - snow, 9 - L a r k remnants.

243

in the lower reaches of the Kolyma river and in north-western Siberia (Fig. 8.10). The isotope variations reveal three peaks of a sharp cooling, when in the lower reaches of the Kolyma, the winter temperature decreased by 9 - 15”C, namely about 33 - 32 ka ago, 27 - 26 ka ago and 22 - 20 ka ago. During the relative warmings that took place 36 - 35 ka and 23.5 - 22.5 ka BP, the climate was nevertheless more severe than to-day, i.e. 6l8O . . .-28 to 30%, against the present 26%, (Vasilchuk et al., 1985). According to Boulton (1979), the maximum development of the Late Wurm glaciation on Spitsbergen took place between 45 and 40 ka BP, when the edge of the Laurentian ice sheet advanced only as far as the region of the Great Lakes. The maximum southern advancement of the Wurm glaciers both in North America and in Europe occurred only about 20 ka BP (Voznyachuk et al., 1978; Chebotareva and Makarycheva, 1974; Dreimanis and Goldthwait, 1973). These asynchronous maximum advances of the glaciers in the northern and southern directions are quite natural and show once again that the time boundary of 22 - 20 ka is important only on a local scale (Fig. 8.1). According to Dansgaard et al. (1982), the last stage of a fast accumulation o f ice by the Greenland ice sheet, as can be seen through borehole Dye3, took place between 40 and 30 ka BP. This time interval, which is characterized by a fast decrease in the sea level and an increase in the ice load in the glaciation centres, is also marked by intense activity of explosive volcanic eruptions on a global scale and oscillations of the geomagnetic field. Thus, the Laschamp - Kargopolovo excursion found in many regions of the world has been dated at between 42 t 5 ka and 33.5 t 5 ka by K/Ar, TL and 14C techniques. Gillot et al. (1979) think that this time interval is characterized by two short individual excursions: the Olby dated at 42 ? 5 ka and the Laschamp dated at 35 s 4 ka. Almost the same age of 38 ka has been obtained by the K/Ar method for a thick horizon of volcanic ash from Campania (y,) described in the Mediterranean deep-sea sequences (Cita et al., 1981; Thunell et al., 1979) and in the loess sequences of the Don basin, particularly in the sections of the Upper Palaeolithic camps in the village of Kostyenki. In this locality, loess containing this ash separates a double “lower humus” layer with forest pollen diagrams, including the archaic Aurignacian stone artifacts with Mousterian traditions, from the “Upper humus” layer with tundra-steppe pollen diagrams and the Solutrean finely flaked stone implements (Praslov and Rogachev, 1982), the latter being also double. Thus, there are grounds to think that the Olby - Laschamp double excursion, 44 - 35 ka, intense volcanic activity, 38 ka, and the beginning of the global cooling, 42 - 40 ka, have a causality and reflect a new, higher level of activity of all the geosphere of the Earth, which was induced under the impact of orbital factors some 42 - 40 ka BP (Fig. 8.1). Oxygen-isotope variations revealed in the ice sheets of Greenland (Dansgaard et al., 1971, 1982) and the Antarctic (Lorius et al., 1979, 1985; Kotlyakov and Gordienko, 1982) as well as in the ground ice of northern Eurasia and in deep-sea cores are indicative of no less than four or five cooling peaks in the time period from 40 to 15 ka (to be more exact these peaks occurred 40, 33 - 32, 27 - 26, 22 - 20 and 18 - 17 ka BP) but none of the appreciable warmings at all (Figs. 8.2 and 8.2). The

244

Wurm interstadials (such as the Arsi, Denekamp, Bryansk, Plum-Point, Novonazimovo and others) have mostly been identified by observing and dating the embryonic buried soils in loess sequences. It seems that the climatostratigraphic importance of these soils as an indicator of the global climatic fluctuations has been overestimated to a great extent. According to the CLIMAP Project (1976), some I8 ka ago the surface sea water temperature was lower than the present one by 6 - 10°C in some North Atlantic regions, by 6 - 4°C in the eastern Pacific and only by 1 - 2°C in the Mediterranean. However, the average summer sea water temperature did not decrease on the whole by more than 2.3"C. Thus, in the Late Wurm the oceans like the continents cooled down mainly in high latitudes, where a permanent pack ice formed, the southern margin of which shifted in the Atlantic as far as SOON Lat. It might be thought that a change of two climatic phases within each kryo-orthoclimathem (namely, at the beginning of the formation of cold maritime climate and close to the disappearance of the continental arid climate - see Figs. 8.2 and 8.8) was induced by the development of extensive fields of permanent sea ice cover in the Northern Hemisphere. The empirical data show that the amplitude of the glacio-eustatic fluctuations in the sea level in post-Eemian time was 120 k 20 m. The oxygen-isotope curves reveal the 6l8O shift from the Late Wurm to the present time of 1.9 k 0.12%0. This value together with the assumed 6l8O of 3 5 % 0 for ice allows one to estimate a change in the sea level at 218 m (Williams et al., 1981). This difference between isotope data

4 20

0

40

l

m

'

I

I/

l

Ill

60 iY

s .

Y)

tl

:

. .

u

~

80

l

/OO

I20

~

140ka

l

vi

~

l

~

l

'

l

~

flla

-. ..

b)

%.

Pacific ocean %,

I

2

3

4 5 a a c d S e

60CT

Fig. 8.11. (a) Comparison of empirically established coasi lines (solid line) and those inferred from oxygen-isotope data on benthic foraminifera (on the basis of 0.01 %o = 10 m ocean layer) (dotted line). (b) Initial isotope curves are given below. (From Williams et al., 1982).

245

and geomorphological evidence (Fig. 8.1 1) has formed a basis for the hypothesis of Pan-Arctic shelf glaciation suggested by a number of scientists, including Mercer (1976), Denton and Hughes (1983), Grosswald (1983) and Williams et al. (1981), Fillon et al. (1983). According to this hypothesis, within isotope stages 5d, 4 and 2 marine glaciation in the Arctic rapidly turned into a shelf glaciation due to snow accumulation at the surface of pack-ice and consequently ice accretion from above. These scientists mentioned that in such a way for only 10 ka in the Central Arctic basin and the Arctic shelf seas a ice sheet of 1.5 to 3.5 km thick was forming, resting against the sea bottom. The isotope effect of this enormous shelf ice sheet of about 8,370,000 km2 in area was quite great, but the eustatic drop in the sea level did not take place (Grosswald et al., 1978; Williams et al., 1981; Denton and Hughes, 1983; Fillon et al., 1983). Theoretically this hypothesis is orderly and attractive, but unfortunately it completely ignores the empirical stratigraphic evidence. According to Volkov et al. (1978), one of the major centres where the Pan-Arctic ice sheet formed was a shelf zone of the Barents and Kara Seas; it was from here that ice advanced into the Eurasian coastal plains up to 65”N Lat. These authors imagined behind the ice front an enormous basin of melt-water running through the Turgai river, the Aral Sea and the Uzboi river into the Caspian Sea and then into the Mediterranean. The degradation of the Pan-Arctic shelf sheet took place during the last glaciation. The last advance of the Barents Sea ice that left behind the Markhida end-moraine belt in the lower reaches of the Pechora river took place already in the Holocene, of which are indicative nine 14C dates ranging from 9.9 to 9.1 k a (Arslanov et al., 1981) obtained for wood taken from sediments underlying this moraine. Grosswald considers these dates a “decisive proof” in favour of the above mentioned hypothesis (Gerasimov and Velichko, 1980; Grosswald, 1983). We think, however, that an extremely local “Markhida belt” is a pseudo-moraine formed as a result of melting of a large assemblage of the gletcher buried ice that persisted there up to the Holocene, possibly for ten thousand years. The notions about the existence of the Kara ice sheet during the Late Wurm are entirely rejected by the presence of the thick syngenetic polygonal ice veins over the shores of the Yamal and Gydan Peninsulas, which continually increase, according to the series of I4C dates (Vasilchuk and Trofimov, 1984), at least from 30 ka to the midHolocene. This is in accordance with the author’s data (Zubakov, 1972a,b) and these of Avdalovich and Bidgiev (1984) on the peats found over the Yamal and Gydan Peninsulas and dated at 40 to 9 ka. The traces of the continuous evolution of organic life in the Arctic Ocean throughout the Pleistocene (Herman, 1975; Clark, 1982; Danilov et al., 1984) are also evidence against the hypothesis of the Pan-Arctic shelf glaciation, which inevitably would have led to the formation of palaeontological hiatus. All this allows us to conclude that the Pan-Arctic shelf glaciation hypothesis has not been supported by the available stratigraphic data and is evidently incorrect. At the same time it should be indicated that the mechanism of the surging deglaciation suggested by this hypothesis is a valuable contribution to the glacial theory and palaeoclimatology.

246

8.4. Spatial reconstruction of the Northern Hemisphere climate during the Late Wiirm, 20 - 17 ka Fig. 8.12 presents a chart of the summer (July- August) air temperature difference for the time of the greatest cooling of the Wiirm glacial, 18 ka BP, and the present epoch. For this reconstruction we have used CLIMAP data (1976) on

Fig. 8.12. Departure of the summer (July- August) air temperature from that at present for 18 ka BP.

Table 8.4. Mean latitudinal temperature differences between the Wurm cooling (18 ka BP) and the present epoch Latitude (day)

AT('C)

80-70

70-60

60-50

50-40

40-30

30-20

20-10

10-0

Mean global temperature

-13.0

-11.5

-8.8

4 . 4

-3.5

-2.3

-1.9

-2.0

- 4.6

241 9.0-55kaB.R

Recent

*I

--*2

18 ku

B.R

a 1 3

Fig. 8.13. Position of the Intertropical Convergence Zone (ITCZ) and circulation patterns over northern Africa during summer (a) and winter (b) at present time, during the Megathermal (9.0 - 5.0 ka BP) and at the Late Wurm (18 ka BP). 1 - monsoon circulation, 2 - trade-wind circulation, 3 area of tropical rain forest. L low pressure; H - high pressure. ~

Fig. 8.14. Deviation of the mean annual precipitation (mm) from present level at 18 ka BP.

~

248

oceanic surface water temperature, 18 ka BP and pollen data for different regions of the Northern Hemisphere (Peterson et al., 1979; Kutzbach and Wright, 1985). In addition, for the USSR territory the reconstructions by Muratova et al., (Burashnikova et al., 1974) and Velichko (1984) have been used. The strongest lowering of summer temperature to 15 -20°C is recorded in northern Canada, Greenland and northern Scandinavia. In western Europe and in the European part of the USSR temperatures lowered by 8 - 1O”C, in the north of the Asian part of the USSR by 8 - 10°C and in the south by 5 - 6°C. In subtropical and tropical latitudes temperature drops did not exceed 2 - 3°C (Table 8.4). Palaeobotanical and pollen data demonstrate drastic changes in the type of vegetation in all latitudinal zones of the Northern Hemisphere. In this case, if in high and middle latitudes forests declined due to temperature lowering, then the shrinkage of the area of the tropical rain forest in Africa and in the Amazon basin was associated with precipitation reduction (Fig. 8.13). Fig. 8.14 presents a chart of differences of annual precipitation sums between the Late Wiirrnian glacial and the modern apoch. To construct this chart, the technique of Khotinsky and Savina (1985) has been used, which suggests the transition from landscape characteristics to quantitative climatic parameters. The natural zonality was reconstructed by pollen data and evidence on the boundaries of deserts and tropical forests and on the lake-level variations dated by the I4C method. As can be seen from the figure, precipitation decreased on almost all continents except for the western and south-western regions of North America, the Mediterranean and the adjacent regions of Afcrica, the Caspian Sea area, northern Mongolia and the Middle East. The highest lake levels are recorded in the Great Basin of the USA (Smith and Street-Perrott, 1982), in the area of the depression of the Great Lakes in Mongolia (Murzayeva et a]., 1984) and in the Middle East Lakes Van and Konya (Murzayeva et al., 1984; Roberts, 1983). In these regions the mean annual rainfall increased by 100- 300 mm. Vice versa, in subtropical regions of Africa precipitation decreased by 300-400 mm and the levels of the Rift Valley lakes and Lake Chad were the lowest. The area of the Sahara expanded, its boundary shifted southwards by several degrees of latitude, the savanna (Sudan - Sahelian) zone vanishing alsmost completely. The tropical rain forest area degraded to a great extent (Hamilton, 1976), being replaced by different types of savanna (Fig. 8.13). Precipitation decreased noticeably in monsoon regions of India (Bryson and Murray, 1977; Swain et al., 1983).

Resume (1) The data of different researchers on different regions show that the time span of such subdivisions as the Tyrrhenian ( = Strornbus), Karangatian, Eernian and Boreal transgressions is greater than that of an interglacial. In other words, there were two or three transgressions with marine fauna of the “Riss - Wurmian” type, which correspond t o isotope stages Se, 7 and 9 according to Kukla (1977) and Bowen (1978) or t o substages 5e, 7a and 7c according to the authors’ data. In the continental sections of the Upper Pleistocene, there are at least two (or three) thermochrons.

However, the U-series and TL dates for them occurring in a range of 240 - 150 and 130 - 90 ka are not yet valid. (2) Tentative temperature reconstructions for the Late Riss - Wurmian optimum (substage 5e) show that the mean global temperature was higher than the present one by 1.6"C in summer, 2.4"C in winter and 2.0"C f o r the year. In high latitude5, the temperature increased in summer by 4 - 7°C and in low latitudes it decreased by 0.5 - 2.0"C. Such temperature patterns led to better moisture conditions on all the continents of the Northern Hemisphere due to increasing monsoon circulation in subtropical regions and convective precipitation in middle and high latitudes. (3) The kryochrons presumably separating the Early and Late Riss - Wurm megathermochrons (the Warthe stage and its analogues) with a probable age of 190 to 140 ka BP have been considered to represent the "Early Wurm" (the Kalinin - Yenisey - Zyryan kryomers). As a result, the last interglacial (the Late Riss - Wurm) appeared under the name of the Karginski - Ponoi thermochron in the schemes of some authors and on the basis of younger 14C dates it waq erroneously placed within the interpleniglacial being correlated with isotope stage 3. (4) Actually the intra-Wurm (interpleniglacial), stage 3, was on the whole colder than the present climatic conditions and subjected to many brief temperature oscillations. In summer the temperature sometimes rose as high as the present one, whereas in winter it was always lower than nowadays. The then climate was much more continental.

This Page Intentionally Left Blank

Chapter 9

CLIMATIC CHANGES THROUGH LATE GLACIAL AND POSTGLACIAL, 16 - 0 KA BP 9.1. Principles of the time classification of the last 16 ka

As one approaches the present time, the bulk of palaeoclimatic information rapidly increases and the accuracy of determining the age of climatic events improves. This makes it possible to list as many climatic events for the last 16 ka as for the previous million years. The major difficulty is to summarize this information on a global basis and t o make “long distance’’ correlation of these events (Fig. 9.1). It is well known that there is a close relationship between vegetation and climate. It is pollen data that served as a basis for constructing a tentative five-term scheme of the Late Glacial - Holocene. Blytt (1882) developed this scheme on peat sections of Denmark and Sweden. Later on this scheme was modified by Nilsson (1964) on the basis of detailed dating (33 I4C dates) of the Agerods-Mosse section in southern Sweden. Neishtadt (1957) and Khotinsky (1977) applied this scheme to the USSR territory, and Auer (1968) and Heusser (1966) to the Southern Hemisphere. The Blytt - Sernander “periods” have been shown to reflect historical development of vegetation associated both with climate change and migrations which lag behind the climatic changes that induce them. In this connection, the time boundaries of pollen zones and “periods” are considerably asynchronous and “running” not only in the meridional but also in the latitudinal direction. Obviously, palynological subdivisions or the boundaries of pollen zones, being only local subdivisions, cannot be used as a chronological basis for constructing global chronostratigraphic scales. Swedish researchers (Deglaciation of Scandinavia, 1979, 1980) attempted to pass to the conventional chronometric scale of the Holocene (similar to the preCambrian) by dividing it into five unified subdivisions with numerical boundaries, 2,500; 5,000; 8,000 and 9,500 years. This seems to be very artificial. Such a division does not promote revealing time structure of the Holocene climatic events and, consequently, the high-resolution long-distance correlation of Holocene sediments. The authors believe that a detailed climate - chronological scale of the Holocene can be based on global climatic events: warmings (thermochrons) and coolings (kryochrons) dated by the radiocarbon method (Borzenkova and Zubakov, 1984). The mass 14C dating of various natural objects, particularly such “sensitive” to temperature change as the northern and upper forest border (Lovelius, 19791, the northern margin of the coral distribution (Taira, 1979), variations in isotopic composition of lake carbonates (Eicher and Siegenthaler, 1976; Morner, 1980; Punning and Raukas, 1985), speleothems (Harmon et al., 1977), ice sheets (Dansgaard et al., 1971, 1982, 1984; Lorius et al., 1985; Kotlyakov and Gordienko, 1982) and underground ice (Vasilchuk et al., 1985), enable us to record the sign change of temperature trend lasting for a hundred years or less in very different facial conditions.

252

J. STAGES AND PHASES OF THE BALTIC SEA

‘

POLLEN ZONES AND VEGETATION (KAJAK A.D. 1976)

MOST IMPORTANT SITES AND THEIR

HEIGHTS ( M )

7 PINUS BETULA LIM

V

1100

-

MUSTJARV

2

1700

-

NOVA

5

PAHILA

6.5

PICEA LIM I V

BETULA

5A 1

LIMNEA SEA ALNUS

LIM I l l

2800

-

3700

-

S6 2

PICEA

LIM I1

LIM I

- 4200

L

561

PUERCUS

OlelKU

11,5

KLOOGA

14.5

ARUSTE

16.6

-

4800

SEA

L Ilb L

AT 2

TlLlA

-15

KOLGA 12.5

RANNANETSA 6.5

LUMANDA

118

6600 - VESIKU L

AT 1

ULMUS

LAKE

.-

KOPU JOELAHTMF

BO 1

PINUS

- 8800 __ ECHENEIS

LAKE

- 9300 YOLOIA SEA

ioaoo

E

Y I1

y

I

BIN

-

N d M N E ICE LAKE G I KEMBA ICE L A K E VOOSE

,3+-

~

BALTIC I C E LAKE ~

16.5

I

- 7200 HASTOGLOIA SEA M - 7600 &NCYLUS

9

IV

L Ill LlTTORlNA

VIHTERPALU

GI

ICE LAKE G I

-12700 __

-8

-

9100

TALLINN KOPU

BETULA

32

3 32

5 4 n ,

SOJMGI 40 U A R A 9 PULL1 11

OPEN VOOOLAND PARKTUNDRA AL

71700 PERIGLACIALE D R 2

r2200

LEMMEOJA

@,a

63

HABER!, MCRlNAKl

32.t

N~MME PALIVEKE

43 46

VETLA,VOOSE

86;e

~

SO

Fig. 9.1. Stratigraphic scheme of the Baltic Sea deposits based o n complex paleontological investigations and I4C datings (Kessel, 1975; Kayak et al., 1976). Abbreviations for diatom groups of different salinities: P + M, poly- and mesohalobous; H, holophilous; I , indifferent species (070) in the bottom sediments of the Baltic Sea. (From Kessel and Raukas, 1979).

253

Fig. 9.2 presents curves of surface air temperature changes for three large regions of the Northern Hemisphere based on the pollen record. These data indicate that the change of the temperature trend in middle and high latitudes of the Northern Hemisphere occurred practically simultaneously, the amplitude of the temperature change in high latitudes (the north of Siberia) being somewhat greater than in middle latitudes. Thus, high latitudes being most sensitive to changes in forcing factors are more representative in revealing global temperature signals (warmings and coolings and changing temperature trends). Moreover, for some regions of lower and middle latitudes the temperature trend can be of the opposite sign (for example, warming in high and middle latitudes is accompanied by cooling in lower latitudes, and vice versa) (Borzenkova et al., 1976; Budyko, 1980, 1984; Vinnikov and Groisman, 1979). Table 9.1 summarizes the chronology of global climatic events over the last 16 ka, obtained by the analysis of a large quantity of actual information and, primarily, of pollen data for north-western Europe (Florschutz et al., 1971; Coope, 1977; Shotton, 1979; Kolstrup, 1980; Lowe and Gray, 1980; Julut et al., 1982; Van der Hammer et al., 1967; Van Gee1 et al., 198011981; Morner, 1980, etc.), for north-western and central regions of the European territory of the USSR (Khotinsky, 1977; Klimanov, 1978; Klimanov and Yelina, 1984; Klimanov et al., 1985; Klimanov and

Fig. 9.2. Reconstruction of summer air temperature ("'2) and total precipitation (mm) (deviations from the present level) over the past 15 ka. (a) North-western Europe, compiled by Borzenkova on palaeobotanical and pollen data of FrenLel (1966, 1967); Julut and co-workers (1971); Iversen (1973); Coope (1977); Shotton (1978); among others; (b) Northern Russian plain and Soviet Baltic Sea area, compiled by Borzenkova o n pollen data of Nikiforova (1982); Klimanov (1982); Klimanov and Elina (1984); Klimanov and Nikiforova (1985); Klimanov and Serebryanaya (1985); Punning and Rauskas (1985); among others; (c) Northern and north-western Siberia compiled by Borzenkova on pollen data of Kind (1974); Nicolyskay (1980); Andreeva and co-workers ( I 981); Klirnanov and Nicoliskaya (1983); and others.

254 Table 9.1. Chronology of climatic events for the last 16 ka. (Compiled by Zubakov after Andreeva et al., 1981; Berglund, 1983; Borzenkova and Zubakov, 1984; Chaline et al., 1985; Dreimanis and Goldwait, 1973; Dreimanis and Raukas, 1975; Geochronologie of the Quat. Period, 1984; Grove, 1979; Heuberger, 1968; Julut et al., 1983; Kaplina and Lozhkin, 1982; Kind, 1974; Klimanov and Nikolskaya, 1983; Khotinsky, 1977; Konigsson, 1984; Nikolskaya, 1980; Patzelt, 1974; Punning and Raukas, 1983, 1985; Schove, 1978; Stotton, 1978, 1983; Zubakov, 1972, 1974). Stages

Substages according to Scandinavian nomenclature

Nannoclimathems - examples of regional events: g.a. = glacial advances, tr = transgressions, rg = regressions, t = thermochrones, k = kryochrons

Subatlantic SA4 (t) “Neoglacial” ~(SA) SA3 (k)

Approximate age of boundaries by 14C

Warming of the XIX - XXth centuries 170(130) “Little ice age in Europe”, g.a. Fernau in the Alps, Goradil tr. of the Caspian Sea 800

SA2 (t)

“Viking warming”, Dyunkerk and Nymphey trs. of the Black Sea, Derbent rg. of the Caspian Sea

SAI (k)

Goshenen g.a. in the Alps, Phanagoriya rg. of the Black Sea

1200

~

~

2500 ____ Subboreal (SB)

SB4 (t) Subboreal warming 2 ____ SB3 (k) Lobben - Zimming g.a. in the Alps

2900? ___ 3500?-

SB2 (t)

Subboreal warming 1

SB1 (k)

Piora - Rotmoos g.a. in the Alps, Raphael g.a. in the Chilean Andes

4600?-

5300 Atlantic (AT) AT5 (t)

~

“Late-Atlantic optimum” t r . Littorina 3 of the Baltic Sea, tr. Fromentain - Nuakshot 6200 Larstig

AT3 (t) _ AT2 (k)

Littorina 2 tr. of the Baltic Sea _ ~ Misox - Frosnitz 1 g.a. in the Alps

AT1 (t)

Littorina 1 tr. in the Baltic Sea, Tyrrell in the Canada

8 0 4 (k)

Frobisher Bay, Cochrane, Cockburn g.a. in North America

8 0 3 (t)

Ancylus I1 tr. of the Baltic Sea

B 0 2 (k)

Gothard g.a. in the Alps, Gousan tr. of the Caspian Sea

B01

“Early Boreal optimum”, Ancylus I tr of the Baltic Sea

-

~

Frosnitz 2 8.a. in the Alps

AT4 (k)

6400 _-

7500

-m E s 4J

P

~ 6800

7900 Boreal (BO)

~

~

~

8300 _ _ 8500

~

8700 __ (t)

9000

~

255 Table 9.1. (coniinued)

-__________--~ ~Nannoclimathems - examples of regional events: g.a. = Approximate age o f boundglacial advances, tr = transgressions, rg = regressions, aries by "C t = thermochrones, k r = kryochrons ~

Stages

Substages according to Scandinavian nomenclature

__.__

-

Preboreal (PB)

PB4 (k)

Daun - Aker g.a. in the Alps, Marresalle (k) of Siberia ___________ Lengholtz g.a. in the Alps, Yabro-Yakha (t) of PB3 (t) Siberia, Yoldia 2 tr. in the Baltic Sea

9300

-~

9600

~

PB2 (k)

Piottino g.a. in the Alps, Pereyaslavl (Russian Plain) - Pit (Siberia) (k)

PB1

Frisland - Polovetsk of the Baltic Sea

__--

(t)

~-

Shimka

(t),

DR3 (k)

--

~

--

~

Yoldia 1 tr. _ _ 10300-

___.____~

Dryas (DR)

9800

--

-

R a - 0 s g.a. (Scandinavian), Valderds (N. America) Ayakli (k) in Siberia ~

11000

~ ~ _ _ . . _ _ _ _ _ _ _ -

AL ( t )

Allerod 11 (0

-

Taimyr - Berelekh

_ _ Middle Dryas - Taberg

-

Two Creeks

___ Neva g.a.

11800

- -

_ _ _ _ _ _ _ _ _ _ _ ~ _ _ 12300

--

DR2 (k) B 0 (t)

Bolling - Windermere Creeks I (t)

-

-

Novomaranka

-

~

Two

_ - 12800 Early Dryas - Luga - Port Huron g.a. __ 13300 RA (t) Raunis Makkinow - Susaka - Kokorevo (t) __________ _ - 13800 - P D (k) Krestsy - Port Brus g.a. - 15300 -Lascaux - Tab-Yakha - Erie (t) LA (t) 16000 ~ _ _ - _ _ _ _ _ _ _ --~

~

DRI (k)

~

~

-

Serebryannaya, 1986 etc.), for Siberia and the north-east of the USSR (Khotinsky, 1977; Zubakov, 1972, 1974; Kind, 1974; Klimanov and Nikolskaya, 1983; Nicolskaya, 1980; Andreeva et al., 1981; Kaplina and Lozhkin, 1982 etc.), for Canada and the USA (Andrews et al., 1981; Colinvaux, 1981; Ritchie et al., 1983; Wright and Porter, 1983; Wright, 1981; Webb, 1975; Webb and Bryson, 1971; Heusser, 1977; Heusser and Heusser, 1980; Heusser and Streeter, 1980; Heusser et al., 1981, 1985; Webb, 1975; Delcourt and Delcourt, 1981, 1983, 1977; Watts, 1980; Adam and West, 1983; Adam et al., 1981; Davis and Jacobson, 1985, etc.). In addition, data on glacier variations in the Alps and Rocky Mountains (Heuberger, 1968; Patzelt, 1974; Kind, 1974; Grove, 1979), on sea level variations (Morner, 1979, 1980; Ruddiman and McIntyre, 1981; Duplessy and Ruddiman, 1984; Mix and Ruddiman, 1984; Ruddiman and Duplessy, 1985; Yang Huai-reh and Xie Zhi-reh,

~

256

1984), and on transgression and regression of the Baltic (Punning and Raukas, 1983) and Caspian Seas (Geochronology of the USSR, 1974; Kvasov, 1975; Varushchenko et al., 1980, History of lakes of the USSR, 1983) have been used. Analysis of these data allows us to reveal about 30 nannoclimathems (NCT). This term means kryomer and thermomer parts of temperature variations lasting from the first hundreds up to a thousand or a thousand and a half years which can be globally observed by 14C dating. Climatic cycles of 1.8 to 2.5 ka revealed by Shnitnikov (1957), Karlstrom (1961) etc. represent, as a rule, groups of nannoclimathems. In Table 9.1, the sequence of nannoclimathems is compared with traditional subdivisions of Scandinavian climatochronological scale serving as nomenot ype.

9.2. On the global temperature trend over the last 16 ka

The curves in Fig. 9.3 show the temperature changes in different latitudinal zones of the Northern Hemisphere during the Late Glacial and Postglacial. The changes in summer air temperature in the latitudinal belt of 60 to 70"N have been obtained by the authors from pollen data for north-western Europe, the northern part of the European territory of the USSR, and Siberia. This curve is compared with independent data for northern Canada (Nicols, 1975), the paleotemperature record from Gotland (Morner, 1980) and with the data of Nesteroff et al. (1983) on the sea surface temperature changes in the southern part of the Mediterranean Sea. The analysis of these curves reveals about 30 global events (warrnings and coolings) lasting from 200 to 2,400 years. Two distinct turning points observed near 9.5 - 9.0 and 5.5 - 5.0 ka BP allow us to divide the climatic history of the last 16- 17 ka into three main units. Following Hafsten (1970) we called them Anathermal, Megathermal and Katathermal. 9.2.1. Anathermal f r o m 16 fo 9 ku BP

This time represents abrupt changes in climate with alternating thermochrons lasting about 600 years, the air temperature being close to today's level, and more shorter kryochrons, with increasing sea ice and readvances of mountain glaciers. The first considerable warming is inferred from pollen data [Interstadial Erie in North America (Dreimanis and Goldwait, 1973), Lascaux in Europe (LeroiGourhan, 1968; Chaline et al., 1985), Tab-jakha in Siberia (Zubakov, 1972)] dates back to 16 t 0.5 ka. During the subsequent warming, 13.2 k 1.O ka, Raunis - Port Stanley? - Kokorev0 (Zubakov, 1972a, 1972b, 1974; Dreimanis and Raukas, 1975) the first appearance of woody plants Betula is recorded on the northern slopes of the Pyrenees and the coast of the Bay of Biscay (Julut et al., 1983). At this time the sea surface temperature in the North Atlantic reached values close to the present ones (Ruddiman and Mclntyre, 1981) and in western regions of North America the summer air temperature inferred by pollen data increased by 3 or 4°C (Heusser, 1977; Heusser et al., 1980).

257

"

I

Fig. 9 . 3 . Deviation of summer air temperature ( " C ) from that at present over the past 16 ka in various geographical belts of the Northern Hemisphere inferred from different proxy geological data. (a) July air temperature in the zone between 60"N and 70"N from pollen and other proxy geological data of Borzenkova and Zubakov (1984). Considerable coolings are correlated with the readvances of glaciers in the Alps (Heuberger. 1968; Grove, 1979; Patzelt, 1972); (b) July temperature estimated from pollen data of north-western Canada, Keewatin and Mackenzie Districts (Nicols, 1975); (c) Palaeotemperature record (58"N) based on oxygen-isotope data on lacustrine carbonate (Chare lime) from Lake Tingstade, Trask Gothland (Morner, 1980). Mean temperature deviation from that at present calculated by Borzenkova from data of Morner; (d) J u l y temperature based on pollen data from Karelia, 62" - 64"N (Klimanov and Elina, 1984) averaged by Borzenkova over four sites; (e) July temperature inferred from pollen data (central part of the Middle Russian Upland), 52" 53"N (Klimanov and Serebryanaya, 1985); (f) Palaeotemperature inferred from a faunal assemblage, southern Mediterranean (Nesteroff et al., 1983). The deviation from the present temperature level was calculated by Borzenkova using data of Nesteroff and other authors. ~

258

During the subsequent interstadials (Bolling and Allerod in north-west Europe, Two Creeks I and I1 in North America) the air temperature rose to the present levels. In the north of the Netherlands and Denmark birch forests appeared; on the British Isles (Coope, 1977) the summar air temperature during the Bolling ( = Windermere) thermochron was close to the modern one. The Allerod warming was the strongest in non-glacial regions of the Northern and Southern Hemispheres. In Siberia it is synchronous with the Taimyr thermochron when woody plants moved to the north. At this time the summer air temperatures in North Siberia were close to, or even somewhat exceeded, the modern ones (Zubakov, 1972; Klimanov and Nickolskaya, 1983; Andreeva et al., 1981). This warming coincided with rapid extinction of mammoths in this region (Vereshchagin, Baryshnikov, 1985). In the Southern Hemisphere, in Chile (at 41"N), the air temperature, as Heusser and Streeter (1980) found, was 4 or 5" higher than the modern one. As the deep-sea data showed, the Northern Hemisphere ice sheets began melting about 15.8 k 0.8 ka BP and sea surface temperatures close to the modern ones were reached at about 13.2 k 0.7 ka BP, i.e. 3 ka earlier than it had been previously believed. According to Ruddiman and McIntyre (1981b), Duplessy et al., 1981; Duplessy and Pujol, (1983), Duplessy and Ruddiman (1984), at least one-third of the ice melted between 16- 13 ka BP. By the end of the Anathermal the Northern Hemisphere ice cover disintegrated by 75%. The deep-sea data showed two-step deglaciation: one between 16- 13 ka and the second between 10-6 ka BP. The melting of the ice sheets stopped between 13 and 10 ka BP. Different proxy data (isotopic evidence, planktonic and pollen data, the ice edge position) show that abrupt climatic changes occurred at that time interval. There is a series of the Dryas coolings, each being 200 to 300 (500?) years long, which resulted in decreasing air and water temperatures up t o the values close to the Wurm cooling maximum. These coolings are most distinct in the North Atlantic and adjoining continental regions. There were at least five coolings: Dryas 1 ( = Luga - List) at about 13.2 ka BP, Fjeros - Neva at about 12.8 ka BP, Dryas 2 ( = Taberg?) at about 12.2 ka BP, Dryas 3 ( = Ra) at about 10.8 ka BP, and (0s) at about 10.5 ka BP. The last cooling, Dryas 3 (Ra-Os, Valders, Ayakli) was the strongest, when the water temperature in the North Atlantic decreased by 6 or 7°C in comparison with the modern, and the ice boundary was close to that of the Wurm maximum cooling. Reconstruction of the surface water temperatures in the North Atlantic for the Late Dryas carried out by Grosswald et al. (1985) by using data of factor analysis of planktonic foraminifera, showed that the strongest water surface temperature lowering up to 7 - 8°C was recorded in the southern part of the North Atlantic, in the Mediterranean and near the western coast of Africa (Fig. 9.4). These estimates are confirmed by independent data obtained by Nesteroff et al. (1983) for the Mediterranean and Ruddiman and Duplessy for the North Atlantic (Duplessy and Pujol, 1983; Duplessy and Ruddiman, 1984; Ruddiman and Duplessy, 1985). In the continental regions adjoining the North Atlantic (northern-western Europe, Great Britain, north-western USSR and the northern Russian plain), northern taiga was replaced by cold steppes and tundra assemblages (Khotinsky, 1977; Punning and Raukas, 1985; Shotton, 1978; Kolstrup, 1980; Lowe and Gray, 1980; Julut et al.,

259

1982). In Siberia, the Late Dryas (the Ayakli kryophase) is characterized by a 5 - 6°C decrease in temperature (Fig. 9. l), and recession of woody vegetation

southward by not less than 500-600 km (Zubakov, 1972a). In northern Siberia, open forest with larch spread up to 57”N and on the Taimyr Peninsula the arctic desert landscape prevailed. The origin of these cold waves revealed against a background of general positive global temperature trend typical of the Late- and Postglacial has recently been the subject of discussions (Flohn, 1979, 1985; Duplessy and Pujol, 1983; Rognon, 1983a,b). The whole session was devoted to this problem at the meeting of the workshop “Palaeoclimatic Research and Models” held in Brussels in 1982 (Ghazi, 1983). The existence of such abrupt climatic events during Late glacial has already been established (Table 9.1 and Figs. 9.2, 9.3 and 9.4). The main difficulty is to determine the time span of these events, which, as different authors believe, make up from 50 - 100 up to 400 - 600 years. If the first estimate is correct, then global temperature changes of not less than 2 - 3°C should have occurred during a few decades. The causes of such abrupt climatic events remain speculative as yet (Flohn, 1974). The most popular hypothesis explaining drastic coolings in the Late Glacial is the

Fig. 9.4. Deviation of the mean annual surface water temperature from that at present for the Northern Atlantic and adjacent areas. Palaeotemperature inferred from faunal (planktic) assemblage and pollen data obtained using transfer functions (after Grosswald and co-workers, 1985). I - ice sheets; a - grounded, b - floating; 2 - sites’ palaeotemperature determination from: a pollens, b - foram; 3 - zone of maximum deposition rate of ice-rafted sands from 25 to 13 ka BP; 5 paths of Arctic icebergs; 6 - contour lines of the negative temperature anomalies. -

-

-

260

surge hypothesis (Flohn, 1979, 1983; Ghazi, 1983; Grosswald et al., 1985). It is assumed that as a result of a rapid increase in sea level due to melting of the continental ice sheets, glacial covers of the coastal zone lying lower than the oceanic level turned out to be in an unstable state. Calving of table icebergs in such regions as the Maine and St Lawrence Bays, the North and Baltic Seas disturbed the equilibrium profile of ice sheets and resulted in the outflow of huge masses of ice to the ocean. As a result, the level of the sea could rise by up to several meters. Actually, submerged ancient coastal lines are the testimony to these drastic changes of sea level (Morner, 1973, 1979/1980). Large numbers of floating icebergs occupying the North Atlantic up to the Bay of Biscay caused a considerable change in the albedo, which could exert certain effects on climatic change (Fig. 9.4). As can be seen from Fig. 9.4, at the end of the Late Dryas 3 the air temperature began rising rapidly in high latitudes: amounting to 5 - 6°C over less than 400 - 500 years. About 10.3 k a BP the air temperature in western Europe and in Siberia was close to the present-day level. This warming is known as the Frisland in northwestern Europe, the Polovetskoe in central regions of Russian Plain (Khotinsky, 1977) and the Isha in West Siberia (Zubakov, 1972). The transition from the Late Dryas kryochron to the Frisland thermochron was conventionally assumed by many researchers to be the Pleistocene - Holocene boundary. At present it is fixed by the choice of boundary stratotype in Gothenburg, Solberga and Moltemyr sections in southern Sweden and is dated through 14C at 10.3 ka (Konigsson, 1984; Olausson et al., 1982). However, this boundary is clearly defined only in the Scandinavian region. In West Siberia the Yabro-Yakha warming, 9.6-9.4 ka (Zubakov, 1972a, 1974) coinciding with the beginning of termination l a in the deep-sea section (Duplessy et al., 1981), can be considered the most clear-cut boundary. The positive temperature trend preserved during the entire Preboreal was interrupted by two comparatively short-term coolings recorded through a readvance of mountain glaciers in the Alps (Piottino and Daun - Aker stages). The highest air temperatures recorded in non-glacial regions of the Northern Hemisphere (Fig. 9.3), correspond to the onset of the Megathermal, the long-term period of warm and relatively stable climate.

9.2.2. Megathermal, 9-5.3 ka BP Although this interval as a whole is characterized by higher summer air temperatures (2 - 3°C above the modern ones in high and middle latitudes), the timescale of its climate is rather complicated. As can be seen from Fig. 9.3, against a background of relatively stable climate there are up to five sufficiently shortperiod but strong warmings separated by short coolings. Two warmings, Lengholtz - Yabro-Yakha, 9.6 - 9.4 ka BP, and Early Boreal (BOl), 9 - 8.8 ka BP refer to the Preboreal and Boreal Blytt - Sernander “periods” and three subsequent (AT,, AT, and AT,) warmings to Atlantic time (Table 9.1). These five “thermal optima” revealed from pollen data correlate with the northernmost position of coral reefs in the Pacific (Taira, 1979).

26 I

During the Megathermal coolings, the Laurentide ice sheet moved. Its remains were preserved on the Ungava Peninsula up to 5.5 ka BP. Also the mountain glaciers began readvancing in the Alps (Misox, Frosnitz 1, Frosnitz 2, Larstig, Piora) and in the Rockey Mountains (Patzel, 1974; Grove, 1979; Heuberger, 1968) (Fig. 9.3 and Table 9.1). Analysis of the Megathermal climatic variations showed that the term “climatic optimum of the Holocene”, widely used in Soviet and foreign literature, is rather uncertain. Gribbin and Lamb (1978) and Imbrie (Absfracfsforthe 9th ZNQUA Congress, Moscow, 1982) have drawn a similar conclusion. As can be seen from Fig. 9.3, the Holocene “climatic optimum” can be understood either as the entire Megathermal, 9 - 5 k a BP or its culminating warming (AT5), 6.2-5.3 ka BP, which the highest oceanic level (0.5- 1.0 m above the modern) and the greatest, for the last 20 ka, expansion of the coral zone in the Pacific occurring at water temperatures higher than 20.5”C correspond to (Taiga, 1979). If the term “climatic optimum” implies the time interval with optimum heat and moisture ratio which provides maximum productivity and maximum plant species variety, then it was not simultaneous in different regions of the globe. Palaeobotanic data show that maximally favourable conditions for developing woody vegetation in the north of Siberia and in north-eastern part of the USSR (Nickolskaya, 1982; Kaplina and Lozhkin, 1982) were recorded during the Early Boreal (BOI) nannothermochron 9-8.8 k a BP. As can be seen from Fig. 9.5, in the maritime lowland of Jakutiya at this time the woody form of birch spread up

Fig. 9.5. Location of sites containing Early Boreal wood fossils on the coasral lowlands of Yakuria and wood fossils (Betula) I4C; 2 wood fossils; 3 and 4 - pollen diagrams; all with Arctic islands. 1 I4C age datings; 5 - present boundaries of the vegetation zones and subzones: AD - Arctic desert, AT - Arctic tundra, T T - typical tundra, FT - forest tundra, NT - northern taiga (Kaplina and Lozhkin, 1982). ~

~

262

to 75”N. As seen from Fig. 9.5, remains are found on the islands Kotelny and Lyakhovsky in the Arctic Ocean (Kaplina and Lozhkin, 1982). In Scandinavia, during the Early Boreal NCT, mountain glaciers were practically absent. In southern Sweden and Norway pine forests developed; this indicates higher summer temperatures and decreasing atmospheric precipitation. In western Norway, broad-leaved forests with elm and oak existed, and summer temperatures were 1 - 2°C above the modern ones (Berglund, 1983). In Karelia and Estonia, the Early Boreal warming clearly manifested itself through the development of thermophills and richer forest vegetation (Punning and Raukas, 1985). In Alaska in the delta of the Mackenzie about 9 ka BP, a close fir forest existed, which then disappeared during the Subboreal and Subatlantic kryochrons (Ritchie et al., 1983). In the Southern Hemisphere during the Early Boreal warming, the highest, over the whole Holocene, summer air temperatures were recorded. In southern regions of Chile (41”N) and in New Zealand, summer air temperatures were 4 - 5°C higher than the modern (Heusser and Streeter, 1980; Salinger, 1981). The Early Boreal thermal optimum was most clearly pronounced in non-glacial regions of the Northern and Southern Hemispheres. This thermal optimum corresponds to the “Hydrological optimum” in tropical and subtropical latitudes. At this time the maximum northward shift of the edge of monsoon rains occurred accompanied by greatly enhanced monsoonal transport of moisture into the tropical continents. In this region one can observe over the last 20 ka the highest lake levels (The Rift Valley, Lake Chad etc.), the maximal discharges of rivers (Niger, Senegal), the expansion of the tropical rain forest belt and a decrease in the desert areas (Sahara, Kalahari, etc.) (Kutzbach, 1983; Kutzbach and Street-Perrott, 1985; Hamilton, 1976; Bryson and Swain, 1981; Swain et al., 1983; Pafaeoecology of Africa, Vol. 4 - 16; Williams and Faure, 1980). Thus, the Early Boreal showed itself not only in the North Atlantic region because of retaining the Laurentide ice sheet still rather large in area and remains of the Scandinavian one in the north of the peninsula. The full deglaciation of these regions took place during the Late Atlantic optimum (ATS) with a minimum volume of land ice and maximum level of the World Ocean. Fig. 9.7 shows the landscape reconstruction of the Northern Hemisphere for the Late Atlantic warming (6.2- 5.3 ka BP). Pollen 14C data were used as the basis for this map. Locations of reference sites are given in Fig. 9.6. For the territory of the USSR the landscape reconstructions by Muratova et al. (1980), Savina and Khotinsky (1982) and Grichuk (1982) have been used. These were supplemented with new data on Taimyr (Andreeva et al., 1981), the north-east of the USSR and Jakutia (Kaplina and Lozhkin, 1982), the north-west of the European USSR (Klimanov et al., 1985) and Soviet Central Asia (Mamedov, 1982; Varushchenko, 1984; Moisture variations . . ., 1980). Landscapes of western Europe were reconstructed by palaeobotanic evidence of Iversen (1973), Frenzel (1966, 1967), Shotton (1980), Van Gee1 et al. (1980, 1981), Van Hammen et al. (1967, 1971), Kolstrup (1980) and Wijmstra (1978) etc. For Canada and Alaska there have been used data by Ritchie et al. (1983), Hopkins et al. (1982), Colinvaux (1981), Andrews et al. (1981, 1984), Nicols (1975), Liu, Kam-Biu (1981), etc., and for the USA, by Barriosky (1981), Heusser (1977), Heusser and Heusser (1980), Delcourt and Delcourt, (1 977, 198l), Muratova

263

and Suyetova (1983), Watts (1980), Webb (1985), Webb et al. (1980), Webb and Bryson (1972), Heusser and Streeter (1980), Winkler et al. (1986), Wright (1961, 1971), Wright and Frey (1965), Wright and Porter (1983), Spaulding (1983), Hevry and Karlstrom (1974) et al. The information for the territory of North Africa was based on the series of publications Pulueoecology of Africa (Vol. 1 - 16), archaeological and palaeobotanical data of Hillaire-Marcell et al. (1983), Petit-Maire (1984), PetitMaire and Riser (1981, 1983), Street (1981), Hamilton’s data on the area of tropical rain forests (1976), deep-sea drilling data from the western coast of Africa (DiesterHaas i976; Rossignol-Strick and Duzer, 1979), data on lake levels variations in the Rift Valley, Lakes Chad and Bosumtwi (Williams and Faure, 1980; Street and Grove, 1979; Gasse, 1980; Servant and Servant-Vildary, 1980; Talbot et al., 1984, etc.).

Fig. 9.6. Location of sites having I4C datings in the Northern Hemisphere used for reconstruction of the Holocene conditions. Solid circles - sites with pollen data (I4C measurements); open circles - Sites where air temperatures were estimated from pollen data using transfer functions of various type,; open squares - sites with deep-sea drilling data; solid squares lake-level data; diamonds sites with archeological data. ~

~

N e CA

Fig. 9.7. Landscape in the time interval of 6 - 5 ka BP. Compiled by Borzenkova from data of Neustadt (1967); Muratova and co-workers (1980); Andreeva and co-workers (1981); Kaplina and Lozhkin (1982); Klimanov (1982); Khotinsky and Savina (1985); among others. 1 - typical tundra; 2 - forest tundra; 3 mountain tundra and shrubland; 4 northern taiga; 5 middle taiga; 6 - dark coniferous mountain vegetation; 7 - southern taiga; 8 mixed broad-leaved/coniferous subtaiga; 9 - broad-leaved forest; 10 - forest-steppe; 1 I - semidesert; 12 - desert. Present limits of the vegetation zones: 13 - southern limit of typical tundra; 14 - southern limit of northern taiga; 15 - northern limit of broad-leaved ~

~

fnvect.

It(

~

n n r t h e r n limif nf wmirlewrf

~

~

The natural zonality reconstructions (Fig. 9.6) were based on about 400 literature sources, 50 of which, mainly later works, were included in the References. Analysis o f the data presented in Fig. 9.1 showed that the most important changes of vegetation occurred in northern (typical tundra and forest-tundra) and in southern (steppe and semidesert) regions. In the north-eastern regions of the USSR the boundary of wood vegetation moved 200 - 300 km northward, the northern taiga (larch and spruce forest) changed for middle and southern taiga (spruce and cedar forest with a small proportion of broad-leaved) (Klimanov, 1982; Nikiforova, 1982). In central regions of the Russian plain, coniferous and birch forests changed for broad-leaved with elm, oak and hazel (Khotinsky, 1977; Klimanov, 1982). In Soviet Central Asia, the forest-tundra zone moved 400 - 500 km northward, typical tundra remaining only in coastal regions adjoining the Arctic Ocean. However, as compared with the Early Boreal warming (Fig. 9 . 9 , climatic conditions during the Late Atlantic warming were more severe; there was no wood vegetation on the coast and on the islands of the Arctic Ocean. In West Siberia, the taiga moved 400 - 500 km northward (Volkova, 1977) and in the region of modern taiga forests on the Ob and Yenisey rivers, fir forests with cedar and elm grew. In the middle part of the Ob, there were found cedar, elm and lime. In the Barabian steppe region, pine and birch forests expanded. In the south of the West Siberian lowland, cold steppe and typical steppe similar to the modern steppe prevailed (Volkova, 1977). In north-western Europe, forest vegetation developed maximally with a great amount of thermophilic forms: oak, hazel, elm and lime (Shotton, 1980; Kolstrup, 1980; Lowe and Gray, 1980). On the North American continent, wood vegetation moved 300-400 km northwards in the central region (in the Middle West), prairies expanded (Webb and Bryson, 1972; Webb et al., 1980). In the south-west, desert and semidesert landscapes similar to the modern ones developed (Smith and StreetPerrott, 1983; Van Devender, 1977, Spaulding, 1983). The area of tropical rain forest in Africa and South America decreased compared with the Early Boreal warming, but exceeded the modern one. The floristic composition of the tropical forests was also richer (Street, 1981; Hamilton, 1976). The area of the Sahara also decreased, giving place to savannas with rich flora and fauna (Sarnthein, 1978; Petit-Maire and Riser, 1983; Van Zinderen Bakker and Coettzee (1972). The landscape reconstruction presented in Fig. 9.7 was the basis for spatial reconstruction of summer air temperature and annual rainfall for the Late Atlantic warming. Summer air temperatures for this time (their departures from the modern) were obtained by the zonal method (Savina and Khotinsky, 1984). This method enables one to reconstruct rather well the shifts of boundaries of natural zones and determine climatic parameters (air temperature and total precipitation) characteristic of certain plant associations. As additional information the air temperature data were used derived on the basis of various transitive functions for different sites (Fig. 9.6). For the USSR territory we used reconstructions of an average July temperature for the time-span of 6 - 5 ka BP carried out by Muratova and Suyetova (Burashnikova et al., 1982), Klimanov (1982) and Khotinsky and Savina (1985). Klimanov’s reconstructions of air temperatures for the European USSR were based

266

on 90 Holocene spectra using the statistical-information method (Klimanov, 1983, 1985) and Khotinsky's air temperature charts for the entire USSR territory were based on nearly 400 Holocene spectra (Khotinsky and Savina, 1985) using the zonal method. Air temperature over the oceans was estimated from the data on sea surface temperature inferred from planktonic fauna of the North Atlantic (Duplessy et al., 1981; Ruddiman and McIntyre, 1984 etc.), the Mediterranean Sea (Nesteroff et al., 1983), the eastern coast of North America (Balsam, 1981) and also from data on coral distribution near the coasts of Taiwan and Japan (Taira, 1979). As can be seen in Fig. 9.8, the strongest warming, to 4"C, was recorded in high latitudes, north of 70 - 75"N. In these latitudes, vegetation and landscape changed much as well. In the USSR European territory, practically all of the typical tundra zone disappeared entirely, remaining only in the Asian part as small areas adjoining the coast of the Arctic Ocean (Taimyr, coastal regions of Jakutiya and Chukotsk). The forest-tundra zone moved northwards by almost 200 - 300 km, corresponding to changes in July temperature of 3-4°C. In middle latitudes temperature differences decreased to 1-2°C and further south they became negative. A small

Fig. 9.8. Deviation of summer (July- August) air temperature from that at present for the Late Atlantic optimum (ATS).

267

Table 9.2. Mean latitudinal differences of summer (July Atlantic optimum (6.2 - 5.3 ka BP) and the present time

~

August) air temperature between the Late _

_

_

_

_

-~ ~ ~

Latitude (day) 80- 70

70 - 60

60- 50

4.0

3.0

1.7

50 - 40

___40-30

~

30 - 20

Mean global temperature ~~

~~

1 .o

0.3

- 0.2

~~

1 .o ~

--

temperature decrease in summer is recorded in Central Sahara, in Central Asia and in Arabia. Apparently these changes are associated with a higher moistening of these regions compared with the modern. Mean latitudinal temperature differences between the Late Atlantic Optimum and the Present time (shown in Table 9.2) were averaged with the consideration of the areas of latitudinal zones. They indicated that the global temperature increase in summer, as compared with the present one, would be 1.0”C. Thus, during the relatively stable and warm climate of the Megathermal the surface air temperature for the Northern Hemisphere as a whole did not exceed 1°C.

9.2.3. Katathermal, 5.3-0 ka BP

As can be seen in Fig. 9.3, the Megathermal-Katathermal boundary can be associated with the first Sub-boreal kryochron, coinciding with the readvance of mountain glaciers in the Alps (Piora stage), 5,200 and 4,700 yr BP and in the Chilean Andes (Raphael stage). On the whole, the last interval of the climatic history of the Holocene, frequently called the Neoglacial, is characterized by a lowering trend of air temperature and increasing instability of climate. As shown in Fig. 9.3, over the last 5 ka five most considerable coolings occurred, when the summer air temperature in high latitudes decreased by 2 - 3”C, and five comparatively small warmings with a 1 - 2°C increase in temperature at high latitudes. The nannokryochrons of the Katathermal (5.3 - 4.8 ka, 3.2, 2.2 - 2.3; 1.5 - 1.6 and 0.7 - 0 ka BP) coincided with glacial advances in the Alps and other mountainous regions. The greatest cooling was recorded at about 2.2-2.3 ka BP and during the “Little Ice Age”, when the Greenland ice sheet started growing again. The last nannothermochron about 1.2 - 1 ka BP, known as the “Viking’s warming”, corresponded t o the Nimphean transgression on the Black Sea and the Derbent regression of the Caspian Sea.

9.3. On possible causes of climate change in the Late Glacial - Holocene Thus, the analysis of data on climate changes in the Late Glacial - Holocene, obtained independently from evidence both for the continents and oceans, shows that climate variations over the last 17 - 18 ka BP were quite complicated (Fairbridge,

268

1983; Borsenkova and Zubakov, 1986). Spatial reconstructions of air temperature for the Northern Hemisphere carried out for two time intervals: the Wurmian cooling maximum (Fig. 8.12 and Table 8.4) and most significant warming of the Holocene - Late Atlantic nannothermochron (Fig. 9.8 and Table 9.2) indicated that the amplitude of global temperature changes in the cycle of glaciation - interglacial would be not less than 4 - 5°C. During the Dryas kryochrons (DR,, DR, and DR,) changes of global temperature by 3 to 4°C could occur over 100 to 300 years and even shorter time intervals.

300

200

100

0

Fig. 9.9. (a) Fluctuations of air temperature at high latitudes (between 60"N and 75"N) since 18 ka BP, solar after Borzenkova and Zubakov. (b) major changes in external and climatic forcing factors: A S radiation as differing (To) from the present level calculated by the authors using data of A. Berger (1978): JJA (June- August) and D J F (December - February), CO, atmospheric CO, (pprnv), averaged by the authors using data of four ice cores from Greenland and Antarctica by W. Berger (1983). (c) estimates of global fall-out of volcanic acids based on p H analyses on Crete ice core (Hammer et al., 1980; Hammer, 1984). -

-

269

Fluctuations in incoming solar radiation due to orbital factors alone are insufficient to explain the complexity of temperature variations in the Late Glacial - Holocene (Fig. 9.9). The effects of astronomical factors were thoroughly studied by Berger (1978). On the basis of these data Pisias has calculated the solar radiation fluxes at the top of the atmosphere on the first date of every two-month interval for the last 30 ka. In part the analysis of these data is presented in Davis (1984). Calculations showed that the pattern of seasonal solar radiation redistribution induced by astronomical factors is more complicated than it was previously assumed. During the last 30 ka in the warm half-year (May-October) there were several maxima in incoming solar radiation: in the beginning of summer (May-June) about 13 ka BP, in the middle of summer (July-August) about 10 - 11 ka BP and at the end of the warm time (September - October) about 5 - 6 ka BP. This redistribution of solar energy should have affected the climatic parameters, in particular the duration of vegetative period and the time of appearance or disappearance of snow cover in high latitudes. I f the first and last maxima of the incoming solar radiation coincided with increasing air temperature in high latitudes (the first with the Raunis NCT, 13.5 - 13.0 ka BP, and the last with the AT, NCT), then the largest flux of incoming solar radiation (11.0- 10.0 ka BP) would coincide with the strongest cooling of the Late Glacial, i.e., the Late Dryas NCT. Although the nature of this cooling is not quite clear, a rapid melting of the continental ice sheets due to high summer solar radiation would be expected to result in considerable albedo changes in the North Atlantic because large volumes of ice would appear in the ocean. One of the factors which can determine the global temperature fluctuations during the Late Glacial - Holocene was short-term ( - 10, - lo3 years) variations of atmospheric CO, content. This problem has arisen only recently when a surprisingly large variations in CO, content were found by the analysis of air bubbles in ice of the Greenland and Antarctic ice sheets (Delmas et al., 1980; Neftel et al., 1982). These data provided the first strong indication that carbon dioxide and its “greenhouse warming” might have been involved in climatic fluctuations in the past. Twenty thousand years ago, during the coldest part of the last glacial period, the CO, content of the atmosphere declined to 180-200 ppm. As the vast ice sheets melted about 16 ka BP, the amount of carbon dioxide began quickly increasing and almost reached the preindustrial level (280 k 5 ppm) at about 10.0 ka BP. The data presented in Fig. 9.9 demonstrated that variations in CO, content under natural conditions during the Late Glacial - Holocene were very complicated. According to Oeschger’s et al. (1985) data, a 50 - 60 ppm change in CO, content can occur over less than 100 years. Recently models have been developed to explain rapid variations of CO, in the atmosphere changing the physical, chemical and especially biological state of the oceans. If the ocean were lifeless, the carbon dioxide would be distributed between water and air solely according to gas chemical solubility. However, in the mixed layer of the ocean rich in microscopic plants and animals, carbon dioxide (in the form of dissolved carbonates) is used by biota to build their tissues and skeletal structures. When they die, carbonate skeletons and organic remains sink into the

270

deep sea, where much of the carbonate dissolves and organic matter oxidizes to insoluble carbonates, which cannot mix with the surface water because of a drastic change in temperature and density. Thus, life in the ocean plays the role of a “biological pump” which delivers carbon from the upper layers to the deep sea. The stronger is this “biological pump”, the less is the CO, partial pressure in the sur-

APmm

t

Fig. 9.10. Relationship between varying air temperature at high latitudes and moisture conditions in various zones of the Northern Hemisphere. (a) Variations in summer air temperature from that at present between 60”N and 75”N after Borzenkova and Zubakov (1984). (b) Moisture conditions in the middle latitudes from studies of the Caspian Sea level fluctuations performed by Zubakov (1974) and Varushchenko and co-workers (1980). 1 - transgression; 2 - regression. (c) Moisture conditions in the subtropical latitudes from pollen data obtained from studies of lacustrine deposits in north-western India, Rajasthan (Swain et al., 1983). Deviations from present rainfall level calculated by the author from data of Swain and co-workers (1983). (d) Moisture conditions in the tropical latitudes from data on the Bosumtwi level variations (Ghana) (after Talbot and co-workers, 1984).

27 1

face water in the atmosphere and vice versa, with lowering of the productivity, the CO, content of the mixed layer in the atmosphere increases. Major factors limiting the productivity are phosphorus and nitrogen coming from the deep sea and whose quantities are great in upwelling regions, where cold bottom waters are formed. If the average productivity in the open ocean makes up 50 g of C/cm2 yr, in the upwelling regions occupying about 0.1070 of the open ocean it may reach as high as 300 g of C/cm2 yr. Along the coastal regions of Peru these values are as large as 1000- 2000 g of C/cm2 yr (McElroy, 1983). All the models trying to explain short-term CO, variations in the atmosphere are associated with productivity changes due to increasing or decreasing upwelling and downwelling processes (Flohn, 1985; Weber and Flohn, 1985), varying circulation rate of deep and surface water, particularly in the Antarctic waters. Siegenthaler and Wenk (1984) assume that comparatively small changes in incoming solar radiation in high latitudes due to astronomical factors (e.g. increased vegetative season), particularly in the Antarctic, are able to induce appreciable variations in productivity of the entire World Ocean and exert an influence on CO, concentration variations in the mixed oceanic layer and atmosphere. Berger assumes the variations in CO, concentration in the Late Glacial to be associated with distinct features of the glaciation process itself, in particular with the surge mechanism of destructing continental ice sheets when during several hundred years enormous volume of meltwater and a great quantity of floating icebergs enter the ocean. The meltwater flow t o the ocean can result in changing the exchange processes between the mixed layer and deep ocean, which to a great extent determines the CO, accumulation in the surface water or, vice versa, its decrease due to rapid carbonate dissolving in the deep sea (Berger, 1983, 1985). As Fig. 9.9 (b) shows, there are three maxima on the curve of changing CO, content of the atmosphere which is the average of four ice cores: about 13 ka BP, between 10 and 9 ka BP and a comparatively small maximum about 6 - 5 ka BP. Minimum values of CO, content are recorded at the time of a cooling maximum about 18- 15 ka BP and after 4.5 ka BP. This change in CO, content agrees satisfactorily with the general trend of surface air temperature variations in high latitudes and with total pattern of change in volumes of continental and sea ice in the Northern Hemisphere. It is interesting t o note that the three maxima in CO, content coincide in time with three maxima in incoming summer solar radiation. The most rapid climate warming between 10 and 9 ka BP that resulted in the Early Boreal temperature optimum (BO I) seems to be attributed not only to increased solar radiation but also to the greatest, over the last 18 ka, CO, content in the atmosphere reaching 350- 380(?) ppm, and even more than 500 ppm according to the Camp Century core (Neftel et al., 1982). As is known from pollen data in the Early Boreal NCT, about 9 ka BP, the highest, for the last 18 ka, air temperatures were recorded in ice-free regions of the Northern and Southern Hemispheres. Using a model of general atmospheric circulation, Kutzbach (1983, 1985; Kutzbach and Otto-Bliesner, 1982; Kutzbach and Guetter, 1984) estimated potential temperature effects of radiation factors (a 7 - 8% increase in summer and decrease in winter of total solar radiation due to orbital factors for the Northern Hemisphere

212

as a whole) in various latitude zones of Eurasia, Africa and Asia. Calculations demonstrated that a considerable increase in summer temperature (up to 2 - 4°C) is revealed in high latitudes of Eurasia; in other regions this warming was relatively insignificant. In winter, temperature lowered a little on all the continents except northern regions of Siberia and the north-east of the USSR, where the temperature increase was not more than 1-2°C. Thus, these noticeable variations in temperature during the Early Boreal NCT cannot be attributed to the radiation factors. Taking into account that the CO, concentration in the atmosphere was considerably higher than at present, we can observe a 6 t o 7°C increase in the summer air temperature inferred from pollen data in the northern regions of the USSR (Taimyr, maritime lowland of Jakutiya, Chukotska) (Andreeva et al., 1981; Kaplina and Lozhkin, 1982; Nikolskaya, 1980; Khotinsky, 1977) and in north of Canada (Ritchie et al., 1983). A sufficiently high CO, content of the atmosphere and large incoming summer solar radiation were responsible for the general positive temperature trend in high latitudes in the Late Glacial and more favourable than modern climatic conditions throughout the entire Megathermal. The onset of general cooling trend of the temperature about 5.0 - 4.5 ka BP coincides with a decrease in summer solar radiation income and a decrease in CO, content of the atmosphere. During the last 3.0-3.5 ka the atmospheric CO, content was 20-30 ppm lower than in preindustrial times (280 & 5 ppm). Although the general trend of air temperature variations in high latitudes over the last 18 ka is synchronous on the whole with solar radiation fluctuations due to astronomical factors and CO, content of the atmosphere, the origin of most considerable coolings in the Late Glacial, Megathermal and Katathermal remains to be cleared up. Along with the surge hypothesis, there is one more mechanism responsible for abrupt climatic change: an explosive volcanic activity. As is known, the climatic effects of volcanism are complicated and ambiguous (Budyko, 1984). The physical mechanism of global temperature variations due to explosive volcanic activity has been well studied on the basis of instrumental data of the surface air temperature over the last 100 years (Asaturov et al., 1986; Rampino et al., 1979; Rampino and Self, 1984; etc.). It was shown that the decreasing transparency of the upper layers of the atmosphere due to the formation of an aerosol layer after a large single explosive volcanic eruption can result in reduction of surface air temperature by 0.5 - 1.0"C (with the thermal inertia of the Earth - atmosphere system taken into account). It is evident that such insignificant and short-term temperature reductions could not leave noticeable tracks in geological records. However, if we suppose that a series of explosive volcanic eruptions could occur during a comparatively short time interval or if we consider longer time intervals (of the order of tens or even hundreds of years), then in both cases the amount of sulphate aerosol ejected into the atmosphere would be much above the amount typical of a single, although very large, eruption. In addition, if thermal inertia smooths, to a considerable extent, the temperature reduction after a single volcanic eruption, then the conservation of increased aerosol concentration in the stratosphere during several years should considerably decrease its effects (Budyko, 1984; Budyko et al., 1986). Taking into ac-

213

count these factors, we can assume that the tracks of temperature effects of volcanic activities at the Earth’s surface should have been preserved in climatological records for the last 18 or 20 ka. Although there does not exist a full chronology of explosive volcanic eruptions over this period, certain studies contain interesting information. The most valuable information on volcanic eruptions in the past was obtained from ice cores (Asaturov et al., 1986; Bray, 1977; Stothers and Rampino, 1983; Hammer et al., 1980 etc.). Large volcanic eruptions inject substantial amounts of silicate microparticles and acid gases into the stratosphere. The polar ice sheets are therefore unique records of past stratospheric aerosol loads over very large latitudinal zones. Such information is recorded in annual layers of higher acidity deposited for several years after the eruption (Hammer et al., 1980; Hammer, 1984). Information about post-glacial volcanism and its climatic impact has been obtained by Hammer et al. (1980) from the Greenland ice sheet evidence for the past 10 ka. According to the data of Taira (1983), three peaks of volcanic activity are revealed in the region of the Aleutian and Japanese islands: about 13.0- 13.5, 1 I .5 - 12 and 10.5 - 10 ka BP, close in time to the Dryas kryochrons (DR,, DR, and DR,). Taira believes that increased volcanic activity and accelerated Pacific plate movements between 13.5 and 10.5 ka BP are associated with the rapid disintegration of continental ice sheets (Laurentide and Scandinavian) which caused a redistribution of the substance mass in the astenosphere and stirring up tectonic processes in the Pacific. A similar hypothesis was earlier proposed by Rampino et al. (1979). The most considerable coolings of the Megathermal and Katathermal can be synchronized with large volcanic eruptions recorded in ice cores from Camp Century. Thus the glacial advance Cochran in the Rocky Mountains about 8.3 - 8.2 ka BP is synchronous with two large explosive eruptions (about 8.2 and 8.1 ka BP); glacial stages Misox and Larstig in the Alps are synchronous with eruptions of the volcano Hekla about 7.4 ka BP, and with the largest (for the Megathermal) eruption of the volcano Mazama about 6.3 ka BP. An unusually large eruption of the volcano Eldgj a (934 AD) is reflected in pollen data from Karelia (Klimanov and Elina, 1984) and north-western Canada (Nicols, 1975) and in oxygen-isotope data of lake carbonates (Morner, 1980) (Fig. 9.3). Geological studies carried out around the volcano El-Chichon after its large eruption in 1982 showed that in the Holocene several large explosive eruptions occurred. Volcanic ash contains a large amount of sulphurous compounds (Tilling et al., 1984). Three layers have been dated by the radiocarbon method. Three datings are available for the first layer: 1.58 f 0.07, 1.6 k 0.2, 1.87 t 0.07 ka BP; the second layer is dated at 1.35 t 0.07 ka BP; and the third has six dates: 0.53 k 0.06, 0.65 k 0.06, 0.57 k 0.06, 0.6 t 0.07, 0.65 t 0.07 and 0.7 k 0.07 ka BP. Comparison of these results with ice core data has indicated that these eruptions are synchronous with two peaks of high acidity signals of the CrCte ice core about 1259 and 623 AD (Hammer et al., 1980). These signals are noted by Hammer and his collaborators as unknown. In our opinion, they are possibly associated with eruptions of the volcano El Chichon. The eruption about 1259 AD, being the cause of a strong cooling at about 1300 AD, is reflected in the pollen record from the north-west of the European part of the USSR (Klimanov and Elina, 1984; Klimanov et aL, 1985).

274

Thus the above data on volcanic eruptions and changes in air temperature according to proxy data show that some cooling of the Late Glacial - Holocene could be associated with the most large volcanic eruptions (or a series of volcanic eruptions) reflected in “acidity” records. Studies of the nature of climatic change throughout the last 18 ka BP demonstrated that the largest warmings and coolings occurred synchronously all over the globe and were associated with: (1) fluctuations in coming solar radiation due to the influence of external or internal forcings determined by astronomical causes or variations in atmospheric transparency because of explosive volcanic eruptions; (2) changes in atmospheric CO, content under purely natural conditions; (3) variations in the albedo of the Earth’s surface and the Earth - atmosphere system due t o changes in sea and land glaciation areas, in the area of desert and arid territories, tropical and middle latitude forests.

The interaction of all these factors in the atmosphere - ocean - kryosphere system, and direct relations and feedbacks between them, have determined the complicated nature of climatic change in this period (Budyko, 1984; Mitchell, 1979). As follows from the analysis of instrumental observations of precipitation, there exists a certain relationship between global temperature variations and moisture conditions in different latitudinal belts. Drozdov (1981) showed that warming in the Arctic area was accompanied by a tendency of decreasing precipitation in middle latitudes, particularly in the regions of insufficient moistening, whereas with coolings a reverse pattern is observed i.e. precipitation increases noticeably in these regions. Lamb (1974), having compared the precipitation anomaly patterns for the global 1930s warming and 1970s cooling, came to the conclusion that precipitation variations are of opposite pattern in different latitudinal zones. With global warming, the moisture conditions are improved in high latitudes and in subtropical and tropical zones of anticyclone (between 0 and 35” Latitude in both hemispheres), whereas precipitation is in deficit on the middle latitude continents. With coolings, vice versa, precipitation increases in middle latitudes and decreases in high and subtropical latitudes. This relationship was confirmed by the author of this chapter in analyzing precipitation anomalies in the 1930s and 1970s using data of 100 longterm instrumental series from the insufficient moisture zone. The results obtained are presented in Budyko (1984). Later on, Wigley et al. (1980) estimating precipitation trends for a warmer earth have constructed charts of the Northern Hemisphere rainfall anomalies for five cold and five warm years chosen over the period from 1924 t o 1975. These charts confirmed the conclusion drawn by Lamb (1974) and the author (Budyko, 1984). Most reliable quantitative estimates of the relationship between changes in global mean air temperature and precipitation in different latitudinal zones have been obtained by Vinnikov and Groisman (1 979), and Kovyneva (1984). Conclusions made in these studies for the f 0.5 range of global mean temperature changes agree well with the empirical data of Lamb, Wigley et al. and the author of this chapter. In

275

the past, as can been from Table 9.3, the range of changes in the global air temperature, for example between the maximum cooling and maximum warming over the last 20 ka, several times exceeded those typical of the period of instrumental observations. It might be supposed that these variations in global temperature should have correlated with considerable changes in the global hydrological cycle and precipitation pattern in different latitudinal zones. Proxy geological data over the Late Glacial - Holocene period obtained by the I4C method represent a rich source of empirical information for studying precipitation patterns in different latitudinal zones with different values of global warmings and coolings.

9.4. Moisture conditions in different latitude zones over the Late Glacial - Holocene: a review of empirical data A lot of information about moisture conditions in the past can be derived from pollen data, lake-level fluctuations, especially closed lakes (lakes without surface outlets), archaelogical evidence, particularly pertaining to arid and semi- arid zones. The largest and best dated empirical information about moisture conditions during the last 20 ka is represented by the lake-level fluctuations of the Afar and Ethiopian Rift (Gasse, 1980; Gasse and Street, 1978; Street and Grove, 1979) for Lake Chad (Maley, 1977, Williams and Faure, 1980; Servant and Servant-Vildary, 1980; Nicolson, 1980), for eastern Mediterranean regions and the Middle and Near East (Bottema and Van Zeist, 1981; Roberts, 1983; Mursaeva et al., 1984; Degens, 1977), for northern and central Mongolia (Vipper et al., 1981), for the Great Basin of the USA (Smith and Street-Perrott, 1983) and pollen data for subtropical regions of Australia (Kerchaw, 1978) and southern Africa (Heine, 1978, 1982; Van Zinderen Bakker, 1972, 1982, 1984; Palaeoecology ofAfrica, vol. 4 - 16, 1971 - 1984). A more sensitive indicator of moisture conditions in the southern area of the middle latitude are the data on the Caspian Sea level fluctuations (Varushchenko et al., 1980). Evidence from archaeological sites of the Mangyshlack, Ustyrt, the south area of Uzbekistan and Kara-Kum and Kysyl-Kum deserts is an excellent source of data about moisture conditions of the Caspian Sea area and Soviet Central Asia over the Holocene (Varushchenko, 1982, 1984).

Table 9.3. The mean latitudinal difference5 of qummer ( J u l y - Augusr) air temperature between the Wurm glacial maximum (18 ka BP) and present time ( I ) and betwecn the Late Atlantic optimum ( 6 - 5 ka BP) and the present time (2) Latitude (day)

90-80 80-70

--______ -13.0 (2) 4.0 (1)

~

70-60

60-50

50-40

-11.5 3.0

-8.5

-5.4

-

1.7

-

1.0

40-30

30-20

20-10

10-0

--_________ -3.5 0.3

-2.3 -0.2

-1.9 -

-2.0

-

~

Global mean _ _ -4.6 1 .o

276

Fig. 9.10 shows a comparison between high-latitude (north of 60"N) air temperature variations, reflecting t o a considerable extent global temperature variations in the Late Glacial-Holocene and moisture fluctuations in middle, subtropical and tropical latitudes. During the Anathermal, global climate warming corresponded to deterioration of moisture conditions in the southern areas of middle latitudes (lowering of the Caspian Sea level) (Moisture variations . . ., 1980) and improvement of moisture conditions in subtropical (monsoon) and tropical regions. The relatively warm and stable climate of the Megathermal is accompanied by the highest lake levels and a considerable rainfall increase in the subtropical and tropical regions. The most considerable warmings and coolings of the Megathermal caused noticeable regressions and transgressions of the Caspian Sea, respectively (Varushchenko, et al., 1980). A reduction of air temperature in high latitudes during the Katathermal induced a rapid lowering of lake levels and amount of precipitation in the subtropical and tropical regions, whereas the Caspian Sea level variations were relatively small. This analysis showed that precipitation in the subtropical and tropical regions is most sensitive to variations in global temperature. Although at present it is obvious that pluvials of tropical and subtropical zone and global temperature increases correlate well, comparatively recently there was an opinion among palaeogeographers that pluvials in these regions (e.g. in Africa) coincide with high latitude global coolings and, vice versa, arid conditions with global warmings. Balout (1955) and then Tricart (1956) were the first to propose the opposite hypothesis about the coincidence of tropical and subtropical Africa pluvials with global warmings and desiccation of these areas with global coolings. Later on, these conclusions were drawn by Fairbridge (1976), Maley (1977), Street (1981) and the author of this chapter (Borzenkova, 1980). 9.4.1. Empirical data on moisture conditions in tropical and subtropical regions between 0 and 25"N and S

The African continent has the best proxy (pollen, palaeohydrological and archaeological) 14C data indicating variations in moisture conditions in subtropical and tropical regions over the last 20 ka. Recent interest in the moisture variations in this region is associated, undoubtedly, with the disastrous drought of the late 1960s in the Sudano - Sahelian zone (Borzenkova, 1980; Nicolson, 1980; Flohn and Nicolson, 1980). A drastic rainfall decrease in this region very sensitive to moisture variations resulted in strong soil erosion and destruction of vegetation, to renew which, as specialists think, it would require several tens of years even with favourable moisture conditions. To solve this problem and to explain possible mechanisms of current droughts in subtropical regions, many specialists recommend to analyse palaeogeographic, archaeological and historical data containing information about rainfall trends for the arid and semi-arid zones (Fairbridge, 1976; Maley, 1977; Nicolson, 1980). Analyzing proxy data, Rognon (1976) came to the conclusion that the Sahara was comparatively moist in the Middle Wurm (40 - 30 ka BP) and desiccation began at about

211

25 or 17 ka BP simultaneously with a global cooling. Maley (1977) made the same inference on the basis of pollen data and lake-level variations of Chad. His data confirm Rognon's conclusion about wetness in the Sahel and its neighbouring regions along the southern margins of the Sahara at about 35 -40 ka BP and during the Boreal and Atlantic warmings (9.0 - 5 .O ka BP) and desiccation of these regions during the glacial maximum 20- 18 ka BP. Probably, during the periods of good moisture conditions southern regions of Sahara were predominantly occupied by savannas, the area of the Sahara, as a desert, greatly decreased, and, vice versa, during the aridity maximum the Sahel zone located on its southern border was trapped by dune sands and did not represent an individual natural zone. According t o Sarnthein (1978), active dune sands occupying now about 10% of the area between 30"N and 30"s during a cooling maximum increased their area to 50%. This led to the formation of two desert regions: one in the Northern and the other in the Southern Hemisphere. Sarnthein believed that desiccation and expansion of the desert belt was associated with a global cooling between 18 - 20 ka BP and the general onset of wetness of the southern Sahara is found only after 12.5 ka BP (Sarnthein, 1978). Sarnthein has also shown that about 9.0-6.0 ka BP the northern border of summer monsoon rains transferring moist air from the Gulf of Guinea shifted to 30"N. At this time according to the data of Petit-Maire and Riser (1981, 1983) the InterTropical Convergence Zone (ITCZ) reached the latitude of the tropics. Extensive lacustrine deposits of 6.5-7.5 ka BP have been discovered between the Erg h e Sakana and the Djebel Takabart (20°30' - 21 ON, 0 - 1OW) in a now hyper-arid area south of Tenezrouft (Petit-Maire and Riser, 1981; Petit-Maire, 1984, 1986). The first lake episode at the southern boundary of the modern Sahara is dated at 9.3 - 7.0 k a BP; the second coincides with the Late Atlantic warming about 5.2 ka BP. The early and mid-Holocene moist phase contributed to the development of the Neolithic culture. The remains of this culture were discovered by Lhote, who found rock drawings in the central regions of the Sahara. Between 9 - 5 ka BP living migration of the African peoples through the Sahara occurred, which shows that the superarid belt was absent at that time (Petit-Maire and Riser, 1983). Pollen data from the Lake Chad deposits, which is a sensitive indicator of moisture variations along the south margin of the Sahara and Sudano-Sahelian zone, confirm the above results. From the data of Maley (1977) it can be seen that the highest levels of Lake Chad were synchronous with global warmings, and the low ones with coolings. According to the data of Servant and Servant-Vildary (1980), the level of Lake Chad began rising at about 15 ka BP and was maximal during the AlIerod warming. The regression of the lake occurred during the Late Dryas cooling about 10.8 - 10.2 ka BP. The highest level and the largest area of the lake took place in the Holocene during the Early Boreal warming (8.9 - 8.8 ka BP). At about 7.0 - 6.9 ka BP there was an interruption in lacustrine deposits and drop of the lake level to values equal to the modern ones. At the same time, moisture conditions deteriorated at the southern boundary of the Sahara desert in north-eastern Mali (Petit-Maire and Riser, 1983). The level of Lake Chad was higher again during the Late Atlantic warming. At the onset of global cooling about 5.0-4.5 ka ago its level began lowering again. It is interesting to note that all lake transgressions

278

in the Late Glacial - Holocene coincide with high levels of Rift Valley lakes and increasing run-off of the Niger river (Gasse, 1980; Pastouret et al., 1978). The data obtained by Degens and Hickey ( 1 974) for tropical Lake Kivu located in the eastern Rift Valley allow us to reveal two pluvials at about 11.8 - 1 1.O ka BP and about 12.3 ka BP; they are synchronous with the global Allerod and Bolling warmings. Interesting results concerning moisture variations in the tropical region of Africa have been obtained by oxygen-isotope analysis of planktonic and benthonic foraminifera of core CH22 KW31 (Pastouret et al., 1978). According these data, the surface and deep water temperatures along the African coast began increasing about 15 ka BP. The maximum temperature of surface water is recorded between 13 - 12 ka BP, which corresponds to the Bolling warming. At this time the greatest volume of freshwater income occurred, which was associated with an increased rainfall in the equatorial region and increased run-off of the Niger river. Some temperature reduction and decrease in freshwater run-off took place between 11.O and 10.0 ka BP, which coincided with the Late Dryas cooling (nCT DR,). Palynological data from Kenya, Uganda, Angola and Tanzania (Van Zinderen Bakker and Coetzee, 1972; Wijmstra, 1978) demonstrated that thermochrons (global warmings) were associated with periods of sufficient moisture and kryochrons (global coolings) with arid phases, accompanied by drastic changes of vegetation type. According to the data of Wijmstra (1978), climate warming during the Atlantic time (9.0 - 5.0 ka BP) was distinctly reflected in pollen data from all regions of tropical Africa. Hamilton (1976) and Street (1981) claim that this time was most favourable for developing tropical vegetation; therefore, the floristic composition was richer and the area of the tropical rain forests greater at that time in equatorial Africa. Thus, during global warmings the equatorial tropical rain forest and savanna zones expand, whereas during global coolings the area of the Sahara desert considerably increases and the tropical rain forests partly degrade (Hamilton, 1976), Moisture conditions at the northern and southern margins of the Namib and Kalahari deserts changed in a similar way. A persistent anticyclone over the Atlantic Ocean at about 30"s and the cold Benguela current between 14 and 10"s are the main reasons of climate aridity in the territory of Namib and in the interior regions of the southern part of Africa. Rainfall at the northern margins of these deserts is determined by monsoons coming from the Gulf of Guinea, whereas precipitation at the southern margin of the Kalahari is associated with the intensity of cyclonic activity at the polar front in middle latitudes. During coolings when climatic zones and circulation systems move to the equator, southern boundaries of the Namib and Kalahari deserts are sufficiently moistened by winter precipitation due to cyclonic activity in middle latitudes, whereas shifting the monsoon to the equator results in precipitation deficit at their northern boundaries. During warmings the pattern is reversed: interior and northern regions of the Kalahari and Namib deserts have a sufficient amount of precipitation while southern regions are in rainfall deficit. Although these regions have been studied little compared with the area around the Sahara, the pollen data confirm this pattern of moisture conditions. Lancaster (1979) showed that during the maximum of-the Wurm cooling (18 - 15

219

ka BP) in the southern regions of Kalahari between 23 and 26"s there was a belt of pans (small dry or ephemeral lakes), which was caused by a phase of heavy precipitation in this territory. It has been assumed that the amount of rainfall in the southern Kalahari at this time was 20- 30% above the present. It is interesting to note that during the 1970s cooling (1972, 1974, 1975) precipitation in this region significantly exceeded the normal and, according to Lancaster, was associated with intensified cyclonic activity in middle latitudes. The second noticeable pluvial at the southern margin of Kalahari was recorded about 11 ka BP, which is concurrent with the Late Dryas cooling in the Northern Hemisphere (Heine, 1978, 1980). Important information about moisture conditions in the south of Africa is obtained from the fossil pollen records of Florisbad (20"40'S), Aliwal North (39"S), Karroo and Cape Provinces (Van Zinderen Bakker, 1982, 1984). During the Postglacial - Holocene at this region two vegetation types alternated: high grass savanna (velds) typical of a colder and wetter climate and dry savanna (karroo) typical of warmings. A dry and warm interval (1 1.6- 1 I .2 ka BP) is most distinctly revealed in pollen diagrams, which could be compared with the Allerod warming in north-western Europe. Later on it was changed to a colder and more humid climate (the Late Dryas in Europe). Most noticeable variations in landscapes occurred about 19 ka BP, when dry vegetation karroo was substituted by wetter savanna (velds). Thus, the analysis of data on moisture conditions in subtropical (monsoon) regions of Africa in the Northern and Southern Hemispheres showed that during global warmings of climate the moisture conditions are improved at the southern boundary of the Sahara and in the northern regions of the Namib and Kalahari deserts whereas in central and southern regions of these deserts precipitation was in deficit. During coolings the precipitation pattern in these regions is reversed as a whole. The bulk of historical evidence, in particular data on the Nile floods, lake-level variations in the Rift Valley on such lakes as Chad, Rudolf, Abja, and other different written sources collected and analysed by Nicolson, showed that such connections between changes in moisture conditions in semi-arid regions of the subtropics of Africa and the temperature have remained during the last 500 years (Nicolson, 1980). For example, during the "Little Ice Age" good moisture conditions were recorded only in the regions adjacent to the Mediterranean coast, whereas along the southern margin of the Sahara and in the neighbouring regions (the Sudano - Sahelian zone) precipitation was in deficit. According to the data of Nicolson (1980), there are three intervals of 7 - 10 years long each during which precipitation decreased sharply in the territories of Sahel and Sudan: 1681 - 1687, 1738-1756 and 1829- 1839. The 1680 drought enveloped the whole Sudano - Sahelian zone and continued over more than seven years. Chronicles retain a record about an unusual event in southern latitudes - snow fall in central Zaire. As the data of Klaus (1980) show, most strong droughts in the Sudano - Sahelian zone recorded in 1639-43, 1681 - 1687, 1711 - 1716, 1738- 1743, 1771 - 1775, 1829- 1834, 1910- 1916, 1944- 1948 and 1968- 1973 are synchronous with the most considerable air temperature reductions in high latitudes discovered in he oxygen-isotope record of ice cores from Camp Century.

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Our data of rainfall trends in the arid and semi-arid zones of Africa and monsoon regions of India for the period of instrumental observations (the last 80- 100 years) demonstrated that the periods with very heavy precipitation in these regions correspond to the Northern Hemisphere warmings (the 1930s and 1950s) and, vice versa, with coolings (the 1960s- 1970s) the rainfall in these regions decreases sharply (Borzenkova, 1980, 1981). Figure 9.11 represents secular precipitation variations at the northern (the Mediterranean Sea coast) and southern boundaries of Sahara (the Sudano-Sahelian zone). It is seen in the Fig. 9.11 that a positive anomaly of precipitation in one of the regions is accompanied by its decrease in the other, and vice versa. A reversed precipitation pattern in these two zones is confirmed by the deep-sea drilling data along the north-western African coast (Diester-Haas, 1976; Rossignol-Strick and Duzer, 1979). It followed from these data that in the latitudinal belt between 15" and 30"N there existed two zones with opposite rainfall trends during the Late Pleistocene- Holocene: one north of 20"N and the other south of 17 - 18"N. During the Late Glacial - Holocene warmings good moisture conditions took place southward of 17- 18"N, whereas is northern regions of the western coast of Africa precipitation was in deficit. The charts of precipitation anomalies in the territory of Africa (the Northern Hemisphere) for the global warming in the 1930- 1940s presented in Fig. 9.12 indicate good agreement between the areas with heavy precipitation and weak I'NU

I0

2U

30

40

30

40 -2

50

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mm I20 80 40 0 -40

-80 -120

I20 80 40

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-160 IN0

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50

60

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80

Fig. 9.11. Secular variations in the mean annual rainfall (deviations from the present level) for two climatic zones of Africa (after Borzenkova, 1980): (A - regions adjacent to the Mediterranean, B Sudan - Sahelian zone. 1 - annual data; 2 - five-year-running data.

28 1

Fig. 9.12. Deviation of rainfall (To) from “normal” level (1891 1970) for current global warming (1931 1940). Shaded areas represent good moisture conditions during the Late Atlantic warming 6 5 ka BP (After Borzenkova, 1980). ~

~

~

precipitation during the current 1930s warming and the Late Atlantic optimum. In the 1960- 70s cooling (Fig. 9.13) and at the end of the last glacial period (18 ka BP) the precipitation pattern in this territory was reverse as a whole. Precipitation deficit (by 5 - 10% and in some regions 15 - 17% below the norm) was recorded along the southern margins of the Sahara and in the Sudano - Sahelian zone. This zone, suffered from the current drought represented during the maximum of cooling 18 ka BP the region occupied by active sand dunes when the Sahara border shifted southward b y not less than 2 or 3” of latitude. It is known that during the drought of the 1960 - 70s it moved southward by a hundred or even more kilometers in certain years, when savanna as an individual natural climatic zone disappeared almost entirely. So, the analysis of empirical data showed that global warmings of different scales that occurred both in the Late Glacial - Holocene and at present induced improved moisture conditions in the monsoon regions of the African continent, whereas coolings caused drastic aridity of these regions. A relationship of this kind can be found in other monsoon regions, in particular in India, Australia and South America. Changes of precipitation in Rajasthan (India) estimated from pollen analysis of Holocene lacustrine deposits of the Lunkaransar and Didwana lakes located at the

282 I0

0

I0

20

30

Fig, 9.13. Deviation of rainfall (To) from “normal” level (1891 - 1970) for current global cooling (1961 - 1970). 1 - regions with heavy precipitation during the maximal Wiirm cooling, 18 ka BP; 2 - margin of sand deserts during the glacial maximum (After Borzenkova, 1980).

margin of the modern Thar desert showed that pluvial phases in this region coincided with warmings and increased monsoon circulation, whereas aridity coincided with coolings and decreased monsoon circulation (Bryson and Swain, 1981; Bryson and Murray, 1977; Swain et al., 1983). As can be seen in Fig. 9.10, an increase in precipitation by 2-3 times, compared with the modern, is recorded in the Megathermal, whereas the desiccation occurred at the end of the last glacial period and during the Late Katathermal cooling. It is interesting to note that the prosperity of the Neolithic culture of hunters in the Sahara and ancient civilization in the Indus Valley (Mohenjo-Daro) occurred simultaneously and coincided with the period of very heavy precipitation in these two regions. Desiccation of the north-western Rajasthan is reflected in the pollen sequence of lake deposits by the appearance of holophyte pollen about 3.7 - 3.5 ka BP. At about 3.5 ka BP sand dunes appeared, which demonstrated a sharp decrease of the rainfall. Desertification of this region seemed to be one of the main causes of the decline of the Harrapa culture in the Indus Valley. As can be seen by pollen data from another tropical region (South America, Eastern Cordillera of Columbia), during the last 14 ka vegetation of some kind of dry grass, paramo, typical of cold and dry episodes was repeatedly replaced with

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vegetation of Andean and sub-Andean forest typical of warm and humid periods (Van Geel and Van Hammen, 1973; Van Hammen, 1985). Fig. 9.14 shows that during the Guantiva interstadial synchronous with the Balling and Allerod warmings in the region of Laguna de Fuquenne (5"N, 2580 m) the Andean forest again invaded the valley, the level of Lake Sabana was rising and air temperatures were close to the modern ones. During the El Abra stadial synchronous with the Late Dryas (DR3) cooling, the forest border was shifted by approximately 800 m, the lake level had drastically lowered and woodland was replaced by grasses (grass paramo). After the El-Abra stadial the climate ameliorated the lake level rose and forests of Andean type invaded this region. According to the pollen record the air temperature during the Megathermal increased by 2 or 3°C. The temperature began to lower at about 5.0 ka BP, then the climate became more favourable and after 3.0 ka BP a new cooling trend resulted in declining forest vegetation. Pollen data from the Amazone river basin, where precipitation is determined mainly by monsoon and equatorial rains, showed that during the Late Wurm and Late Glacial coolings the tropical rain forests are partly degraded to small regions,

Ndrih-west

Europe

e

Colombia, Laguna de Fuquene Pollen Vegetali,

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Fig. 9.14. Vegetation and climate over the last 20 ka in the tropical region of Colombia (Laquna de Fuquene, 2580 m). Correlation with climatic events in western Europe and subtropical Africa (After van Geel and Van der Hammen, 1973; Servant and Servant-Vildrary, 1980).

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refuges (Street, 1981; Padug and Quittgo, 1982), being replaced by savannas of different types. Thus, data on vegetation and climate change in the tropical regions of Africa and South America agree quite well. A dry climate with savanna vegetation type prevailed in these regions between 15 and 12 ka and after 3 ka BP. During the Early Boreal and Holocene warming a considerable increase of rainfall resulted in the replacement of open areas with shrubs by tropical rain forests. A similar moisture pattern is also observed on the Australian continent, where northern regions are influenced by monsoon circulation. During the cooling maximum between 17 and 20 ka BP in north-eastern Queensland, where precipitation falls during summer monsoon rains, its amount decreased by 2 or 3 times compared with the modern (Kerchaw, 1978). Sclerophyll taxa prevailing there demonstrate strong climate aridity. In southern and south-eastern regions, aeolian activity increased and sand dunes became active at this time. Lake Leake dried up between 25.0 and 15.0 ka BP and lake levels in southern and south-eastern Australia were lowered (Kerchaw, 1978). Moisture conditions in north-eastern regions began improving after 15 ka BP, when in pollen diagrams from lakes Eramo and Linch's Crater the percentage of angiospermous taxa increases and of sclerophyll taxa decreases. Between 14.0 and 12.0 ka BP precipitation around Lake Torrence increased and vegetation became more diverse, whereas, according to the lacustrine deposit record of Lake Frome, sand dunes had been preserved there to 13 ka BP. The Late Atlantic warming of the Holocene was accompanied by a considerable improvement of moisture conditions in northern and north-eastern regions, where sclerophyll taxa disappeared being replaced at first by mesophytic forest and then by vine tropical forest, existing under higher rainfall. At that time precipitation around Lynch's Crater was 1000- 1500 mm above the modern (Kerchaw, 1978). In south-eastern south Australia the wettest period of the last 10 ka occurred between 6.9 and 5.0 ka BP. At this time Lake Leake and other lakes in southern and southern-eastern Australia were of a very high lake-level status. The shift to the south of the summer monsoon rain edge when the climate warms can be observed through changes in vegetation type around Lake Frome located at the boundary of summer and winter precipitation. (Zubakov and Borsenkova, 1984). During the Early Boreal warming between 9.0 and 8.0 ka BP, arboreal vegetation prevailed in the region around Lake Frome; sufficient moisture conditions had been preserved to 4.0 ka BP; at about 7.0 ka BP precipitation somewhat decreased, which was synchronous with the increasing aridity in subtropical regions of Africa about 6.9 ka BP (Petit-Maire and Riser, 1983). Thus, the above empirical data on changing moisture conditions in the tropical and subtropical (monsoon) regions of the Northern and Southern Hemispheres revealed a close and unique relationship between rainfall at these latitudes and global coolings and warmings over the last 20 ka. Empirical data on lake-level variations during the Wurm cooling maximum (18 ka BP) and the Late Atlantic warming (6 ka BP) presented in Figs. 9.15 (a, b) confirm the inference that monsoon region pluvials correspond to global warmings and arid phases to global coolings. As can be seen from this figure, northwards, between 25 - 40"N and S, there are areas with reversed moisture trend. They are the northern

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coast of Africa adjoining the Mediterranean, the eastern Mediterranean, Middle East, the northern part of Arabia, the depression of the Great Lakes Basin in Mongolia, pluvial lakes of the Great Basin in the south-west of the USA. In the Southern Hemisphere this area includes the greater part of southern Africa and South America.

Fig. 9.15. Lake-level status: A - 6 - 5 ka BP, B - 18 ka BP. Compiled by the author from data of Zubakov (1974); Street and Grove (1979); Varushchenko and co-workers (1980); Magny and Olive (1981); Street-Perrot and Roberts (1982); Smith and Street-Perrot (1983); Khrustalev and Chernousov (1983); History of lakes in USSR (1983).

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9.4.2. Empirical data on moisture trends during the Late Glacial - Holocene between 25 and 40”N and S

The most reliable empirical data obtained by I4C are available for the southernwestern regions of the USA. The reconstruction of pluvial lake levels of the Great Basin accomplished by Smith and Street-Perrott (1 983) for different time intervals of the Late Glacial - Holocene showed that between 18.0 and 16.0 ka BP lakes had high levels. As warming developed, the lake levels fell, the lowest levels being recorded during the Early Boreal and Atlantic warmings (9.0- 5.0 ka BP). By that time deserts and semi-deserts of modern type had formed in the western United States. Adam and West’s data (1983) obtained from pollen records from Lake Clear (California), showed that during the cooling maximum (18 ka BP) the air temperature decreased in this region by 7 or 8°C and precipitation increased by 300 or 350% compared with the modern. At the Atlantic time it became warmer and drier, and precipitation decreased by 200 or 250 mm. In analyzing DSDP data from the north-western coast of Africa it can be concluded that moisture conditions in this region improved during global coolings, and aridity increased during warmings (Diester-Haas, 1976; Rossignol-Strick and Duzer, 1979). This has been confirmed by the data of Rognon et al. (1984) on the presence of buried soil in Morocco dated at 20.0- 18.0 ka BP, i.e., the time of the Wiirm cooling maximum. The conclusion about increasing precipitation on the northern coast of Africa in the regions adjoining the Mediterranean Sea is drawn also for the modern coolings of the 1970s (Fig. 9.13). It is difficult to analyze the moisture conditions of the eastern regions of the Mediterranean because 14C dated evidence is not abundant. Lake-level variations together with pollen data are the best indicators of climate in these regions (Horowitz, 1979). Direct physical methods, in particular determining the oxygenisotopic composition of natural ground water, freshwater (lake) carbonates and speleothems, also give very interesting results. Recently Magaritz and Heller (1980) proposed a new method for reconstructing climatic conditions in arid regions by using oxygen-isotopic analysis of shell carbonate of the fresh-water mollusc (Levanita

caesareana).

As known, modern L. caesareana shells from an arid zone (precipitation < 300 mm) are enriched, by approximately 2%0, with heavy oxygen (l8O) compared with shells from a temperate climatic zone. In this study shells of L. caesareana found in archeological excavations of two prehistoric caves in Israel were used. The reconstruction of moisture conditions in this area over the last 1 5 ka showed that sufficient moisture between 14 and 12 ka BP was changed by a drastic drying up of climate between 10.0 and 10.5 ka BP (the Late Dryas cooling in Europe), after which moisture conditions improved again, with a humid “period” culminating at 9.0-5.0 ka BP (Magaritz and Heller, 1980). Palynological data by Saad for the western Egypt desert showed that sufficient moisture conditions existed there between 18.0- 13.0 ka and 9.0- 6.0 ka BP. After 6.0 ka BP the climate remained warm but more arid (Saad, 1979). Moore (1979) obtained interesting data on climate in the Euphrates Valley when analyzing archeological evidence. The best moisture conditions in this territory are

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recorded during the Atlantic warming (between 9.0 and 6.0 ka BP), and during the second part of Subboreal time (after 3.0 ka BP), the climate became severely barren and extremely dry. Empirical data on changing moisture conditions in the eastern and Mediterranean regions are based on lake-level variations from western Iran and eastern and southwestern Turkey (Van Zeist and Woldring, 1978; Bottema and Van Zeist, 1981; Degens and Kurtman, 1977; Roberts, 1983; Street-Perrot and Roberts, 1983). Although correlation of pollen diagrams from these regions is rather complicated, however, the sequence of pollen records indicates that the vegetation developed there mainly in similar ways. In the early Holocene open vegetation was spread wider in these regions than at present and resembled modern desert-steppe vegetation of the interior regions of Iran and Afganistan. During the Late Wiirm and Late Glacial coolings, the percentage of arboreal increased noticeably. Lakes Van and Konya had the highest levels during the Wiirm cooling maximum (23.0- 18.0 ka BP) and the Late Dryas cooling (1 1.0- 10.0 ka BP). The level of Lake Van at about 18.0 ka BP exceeded the modern one by 1.5 times and the maximum regressive phase coincided with the Early Boreal warming when the lake level dropped by 340 m and the water volume decreased by 12 times (Degens and Kempe, 1977). Moisture conditions began improving and the levels of Lakes Van and Konja began rising after 4.0 ka BP concurrently with the global cooling trend. As palaeogeographical studies of lacustrine deposits in northern Mongolia (the depression of the Great Lakes) have shown, lake transgression and the expansion of arboreal vegetation occurred at the beginning and in the second half of glacial stages when the annual precipitation in this region increased to 400 - 700 mm compared with 100 - 300 mm at present (Murzaeva et al., 1984). At the beginning of the Holocene at about 9.5 ka BP, a sharp lake regression occurred and woodland was substituted by steppe and semi-desert. However, during the Megathermal in central Mongolia, lakes were again filled and in the mountainous regions forests with birch, larch, cedar and fir rapidly developed (Vipper et al., 1981). Probably, moisture conditions in Central and North Mongolia improved during the Wiirm cooling and Atlantic warming due to different reasons. Improvement of moisture conditions during global coolings depended on increasing winter precipitation due to stronger cyclonic activity, whereas during warmings it was determined mainly by local convective precipitation and, possibly, during the most considerable Megathermal warmings by northward penetration of the monsoon rains. It could not be ruled out that the latter caused a noticeable improvement of moisture conditions in Central Asia and the Caspian Sea area during the Early Boreal and Atlantic warmings. According to Mamedov (1982) and Varushchenko (1984), in the territory of modern deserts Kara-kum, Kyzyl-Kum and in the desert regions of Mangyshlak and Ustyurt (the Caspian Sea area) between 9.0 and 4.0 ka BP rainfall was considerably above the modern (by 200-300 mm), which is reflected in lacustrine deposits, buried soils of this age and archeological evidence about the occupation of these areas by ancient man at about 8.0 - 7.0 ka BP. Global cooling starting at about 4.5 -4.0 ka BP resulted in a noticeable decrease of precipitation and desiccation of this region. A comparatively brief interval of good

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moisture conditions was recorded in Ustyurt and southern Turkmenia during the “Viking warming” at about 1000- 1200 AD. However, the “Little Ice Age” cooling led again t o desiccation and desertification of this area (Varushchenko, 1984; Abramova, Turmanina, 1982).

9.4.3. Empirical data on moisture condition variations in middle (between 40 - 45 ” and 60 “Nand S) and high (north of 60”N) latitudes

As is known, the greater part of the landmass of the Northern Hemisphere is located in this latitudinal belt. Pollen data and lake-level fluctuations are one of the most abundant sources of paleohydrological and palaeoclimatic information for mid- and high-latitude continental areas. Such large lakes as the Caspian Sea, Aral, Leman, Balkhash, Sevan etc. are perfect indicators of moisture variations in central Eurasia. The investigations concerning lake-level fluctuations over the past 20 ka show that main palaeohydrological events occurred more or less synchronously in all central Eurasia. The Caspian Sea level variations determined for 80% by the Volga run-off are a sensitive indicator of changing moisture in the central regions of the European USSR. Data on variations in the Caspian Sea level obtained by the 14C method and presented in Fig. 9.10 confirm the synchronism between the lake transgressions and global coolings as well as between regressions and warmings. The most considerable regression of the Caspian, the Kulalian, (9.1 - 8.9 ka BP), is concurrent with the Early Boreal optimum. Two others, the Zhelandy (7.2 - 6.5 ka BP) and the Izberbash (about 5.9 ka BP) correspond t o the Megathermal thermal optima. Magny and Olive’s (1981/1983) data show that abnormally low levels of Lake Leman were recorded also during the Early Boreal warming and the Megathermal. The cooling that began is Subboreal time resulted in raising the level of Lake Leman. The Holocene history of Lake Balkhash (Khrustalev and Chernousov, 1983) tells us that the most significant lake transgressions coincide with global coolings and regressions with warmings. The Early Boreal warming was accompanied with a decreased lake level at about 8.3 f 0.1 ka BP, and the late Dryas cooling coincided with the ancient Balkhash transgression at about 10.3 f 0.2 ka BP. The reconstruction of precipitation during the early Boreal warming by Khotinsky and Savina (1984) for the USSR indicates significant climate aridity in the European USSR and Western Siberia. Analysis of palaeobotanical data on north-western Europe (Punning and Raukas, 1985) also shows a considerable precipitation decrease on the European continent, including the territory of Scandinavia, where at this time mountain glaciers occupied a very small area or were absent at all. During the Early Boreal warming, there occurred drastic drops of pluvial lake levels of the Great Basin in the USA (Smith and Street-Perrott, 1983), formation of deserts and semi-deserts of modern type in the south-west of the USA and maximal expansion of prairies in central regions of the North American continent. Similar changes of the rainfall pattern are observed also in middle latitudes of the Southern Hemishere. Precipitation data derived from the pollen content of sediments in the Taiquemo and Alerce lake basins (41 - 56”S, Southern Chile) decreased by about 400 - 500 mm below the present one.

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Only in high latitudes (north of 65"N) in the north of Asia (Taimyr, coastal lowlands of Jakutia, Chukotka Peninsula) (Andreeva et al., 1981; Kaplina and Lozhkin, 1982) and North America (Alaska, Canada) (Heusser et al., 1985; Andrews et al., 1981) palaeobotanical data show a considerable increase of rainfall (100- 150 mm per year above the present one) during the Early Boreal optimum. So, analysis of empirical data showed that during the Early Boreal warming (BO,), precipitation patterns were very different in various latitude zones. A sharp drying of the climate in middle latitudes was accompanied by a significant increase in precipitation in monsoon regions, which caused a very high level of tropical and subtropical lakes in both hemispheres over the last 20 ka (Fig. 9.15). Further development of global warming in the Megathermal produced some attenuation of the monsoon circulation intensity and a decrease in precipitation in these regions, while in certain regions of middle latitudes, in particular on the European continent, moisture conditions improved. Fig. 9.16 presents the reconstruction of annual rainfall (in departures from the modern) for the Late Atlantic warming (AT5) (6.2 - 5.3 ka BP). The palynological evidence and lake-level data mentioned above were the basis for compiling these

Fig. 9.16. Deviation of the annual mean total precipitation (mm) from that at present for the Late M a n tic warming (ATS).

290

charts. For the transition from vegetation type data to quantitative estimates of climate (precipitation) the zonal method by Khotinsky and Savina (1982) has been used. The earlier reconstructions of precipitation for the USSR territory for the time 6.0-5.0 ka BP (Klimanov, 1982; Burashnikova et al., 1982) have been supplemented with new results obtained from palynological data on the north-west and north-east of the USSR (Punning and Raukas, 1985; Klimanov and Nikolskaya, 1983; Nikiforova, 1983). As can be seen from Fig. 9.16, the most considerable decrease of precipitation during the late Atlantic warming was recorded in central regions of North America, whereas on the European continent moisture conditions were close to the modern. Moisture conditions improved much in monsoon regions and high latitudes, north of 60"N. Comparable analysis of precipitation reconstruction charts and lake-level variations for the Wiirm cooling maximum (Figs. 8.14 and 9.15b) and for the Late Atlantic warming (Fig. 9.16 and 9.15) points out that the relationship between global temperature and precipitation in various latitudinal zones obtained from modern data is confirmed by paleoclimatic information although the amplitude of these variations in the past, as should be expected, was considerably greater. This can be seen from Fig. 9.17 presenting relative values of precipitation anomalies averaged over latitudinal zones for these two time intervals. For comparison, in the graph there are curves representing precipitation variations during the modern warming (the 1930s) and cooling (the 1970s).

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Fig. 9.17. Mean latitude anomalies of rainfall (To) from the present value for two levels of global warming (current, the 1930s and the Late Atlantic, 6 - 5 ka BP) and for two levels of global cooling (current, 1960s and the Late Wurm, 18 ka BP). Calculated from Figs. 8.14 and 9.16. Current data of Borzenkova (Budyko, 1984).

Table 9.4. Variations in precipitation in different latitudinal zones depending on changes of global temperature in the cycle of glacial and interglacial (for the last 20 ka) ~-

Tropical and subtropical 25" -45" 45" - 60" latitudes 0 - 25" -~ _ __ __ __ - _ _ _ _ _ - _ _ _ _ _ - _ - _ ~ - _ _ _ _ Weakening of monsoon circulaInsignificant improvement The role of cyclonic preof moisture conditions tion. A southward shift of the cipitation increases and moisture condithat of convective dezone, expansion of desert zone, creases. drop of late levels.

Maximal area of tropical forests and minimal area of deserts. Northernmost position of ITCZ, increasing intensity of monsoon rains. Maximally high lake levels

High latitudes (north of 60") _____ Decrease in precipitation, increase in zonal transfer and decrease in meridional transfer.

Improvement of moisture conditions due to meridional circulation. Maximal development of taiga handscapes, development of forest soils. Movement of arboreal vegetation for northward. _____ Gradual improvement of moisture Gradual lowering of lake improvement of moisture Improvement of moisture condiconditions, increasing monsoon levels during warming, dis- conditions due to precipitions during warmings o f Postcirculation, elevation of lake level placement of cyclonic acti- pitation of cyclonic origin. glacial. in the Rift Valley of Africa along vity in northern regions Processes of soil formathe southern margins of the tion. Elevation of the Caspian sea level. Sahara. _________-__ _ _ ----__Sharp decrease in precipitation, Strong aridity. Development of High lake levels in the Aridity, development of maximal expansion of desert area, Great Basin (USA), Middle soil fluctuation and loess kryogenesis processes, maximal partial degradation of the tropical East, Eastern Mediterraarea of sheet glaciation accumulation in Europe rain forest nean regions, Northern and Asia. Strong winds, Africa due to precipitation dust storms. of cyclonic origin and sharp decrease of ebaporation. Aridity in interior regionx o f Asia. Regression phase of lakes of the Great Basin, in the Middle East, interior Mongolia particularly, at the beginning of warming

Rainfall close to the modern. Maximum of soil formation in Europe and Central Asia. Regression of the Caspian Sea.

292

As can be seen from Fig. 9.17 and Table 9.4, there is a certain relationship between variations in global thermal and moisture conditions for large territories in different latitudinal zones. At the same time, local relationships between temperature and precipitation are rather complicated and, to a considerable extent, ambiguous, which is determined mainly by microclimatic conditions. Such indirect indicators as lake-level variations, soil and landscape types characterize moistening of vast areas and give an idea about macrocirculation processes within a certain climatic or latitudinal zone. Most clear relationships between global temperature variations and moisture conditions are found in tropical and subtropical regions where precipitation depends on the intensity of monsoon circulation and the position of the ITCZ. During all the interglacials and interstadials the intensity of monsoon circulation increased, which is associated with distinct features of seasonal redistribution of solar radiation due to astronomical factors (Rossignol-Strick, 1985; Kutzbach and Street-Perrott, 1985). The warmest isotope stages (5e, 7, 9, 11) were accompanied by most favourable moisture conditions over the greater part of the Sahara, in Arabia, and in monsoon regions of India and Australia. When coolings are comparatively small like the “Little Ice Age”, Subatlantic and Subboreal, moisture conditions in the middle latitudes improve mainly on account of increasing zonal transfer. However, with stronger coolings (DR,, DR,, DR, and Late Wurm), when the general moisture content in the atmosphere decreased considerably, a precipitation deficit occurred in middle latitudes. In these regions loess accumulation was promoted also by increasing wind activity (Table 9.4). Moisture conditions improved at this time only within a narrow zone of subtropical latitudes (Great Basin, USA; Middle East; the Mediterranean coast of Africa), where precipitation mainly depends on cyclonic activity at the polar front of middle latitudes. Now we shall consider possible mechanisms of the relationship between global temperature change and precipitation distributions over the latitudinal zones. One of the mechanisms explaining the relationship between temperature and precipitation is the considerable change of the average meridional temperature gradient determining the transfer of moisture to the interior regions (Budyko, 1974; Drozdov, 1983a, b). However, this is only one of the mechanisms that causes the redistribution of precipitation by latitudes with changing global temperature. Recently the concept of possible shift of climatic zones to the pole or equator due to global warming or cooling (Budyko, Vinnikov, 1979; Bryson and Murray, 1977; Lamb, 1974) has been widely used. It can be assumed that the realization of this scheme is carried out through a corresponding replacement of the main pressure centres of action in the atmosphere, which are reflected in empirical data. Borzenkova (1981) has shown that during global warmings the Azores High moves northwards and during coolings it returns southwards to the equator. Bryson and Murray (1977) believe that these movements of the Azores High induce variations of the Inter-Tropical Convergence Zone (ITCZ) whose secular and seasonal locations determine precipitation patterns in subtropical and monsoon regions of Africa, India, South America and Australia. Having analyzed causes of the current Sahelian drought, Bryson showed that comparatively small variations of global

293

temperature, only by 0.3"C, lead to noticeable equatorwards shifts of the Azores High and ITCZ by almost a degree of latitude. This is enough to cause considerable changes in rainfall in subtropical regions of Africa and India. Recently Kovyneva (1984) has carried out a statistical analysis of modern variations in surface air temperature and pressure fields. The results obtained by using the available archive of instrumental data of atmospheric pressure and air temperature for the entire Northern Hemisphere demonstrated that with a 0.5"C warming of the hemisphere the cyclonic activity in the 55 - 70"N zone weakened in the cold season and the Siberian High becomes stronger shifting somewhat north-west. The centre of the Azores High shifts north-east, and pressure grows in the Icelandic Low, whereas the Aleutian Low, vice versa, deepens. As a whole, the cyclonic activity becomes more intense over the Pacific and somewhat weakens over the Atlantic. With global cooling the character of pressure variations over the indicated regions is the reverse as a whole. It can be supposed that if comparatively small global temperature variations occurred in the period of instrumental observations they can cause such noticeable changes in the location of main centres of action and influence their intensity, then the more considerable global temperature variations that took place during the last 18 ka BP should have exerted more pronounced effects on precipitation, particularly in the regions of insufficient moisture as the most sensitive ones. One of the possible mechanisms of changing precipitation in tropical and subtropical Africa is associated with shifting of the main centres of action (in particular, the Azores High), which determine the location of the Inter-Tropical Convergence Zone. With global warmings it shifts farther northwards, and summer monsoon from the Gulf of Guinea penetrating to the north would promote a high precipitation along the southern margins of the Sahara and its surrounding regions. During global coolings the Inter-Tropical Convergence Zone migrated to the equator and the southern regions of the Sahara and Sahel are in the zone influenced by dry trade winds. It can be thought that a similar mechanism was responsible in the past for the precipitation pattern in the monsoon regions of Africa, India, southern America and Australia. At the same time, with shifting climatic and circulation zones, the precipitation variations in tropical and subtropical regions are determined by increasing (during warmings) or weakening (during coolings) intensity of monsoon circulation depending on the temperature difference between the continent and ocean. Kutzbach (1983, 1985), Kutzbach and Street-Perrott (1985) and Rossignol-Strick (1985) have noticed that the tropical and subtropical pluvials correspond to increasing summer and decreasing winter solar radiation due t o astronomical factors. Such changes in solar radiation incidence promote an increase of the intensity of the summer monsoon circulation. Melting continental glacial sheets and the flow of great masses of fresh and comparatively cold water t o the ocean seemed to maintain noticeable temperature differences between the ocean and the continent during the deglaciation. Fig. 8.13 represents summer and winter locations of the Inter-Tropical Convergence Zone at the present time (Misserli, 1980), during the early Boreal and Atlantic warmings (9.0- 5.0 ka BP) and during maximum cooling (18 ka BP). These

294

reconstructions made from paleobotanical (Hamilton, 1976) and paleogeographical data (Van Zinderen Bakker, 1982, 1984) show considerable changes in the tropical rain forest and the Sahara's area over the last 18.0-20.0 ka BP. As can be seen from Fig. 8.13, the winter location of the ITCZ determines the northern border of the tropical rain forest, which would grow only under conditions of uniform rainfall over the entire year, and the summer location - the southern border of the Sahara (or the northern edge of the Sudano-Sahelian zone). A shift of the ITCZ northwards (approximately by 1 - 2" latitude) during the Atlantic warmings of the Holocene determined the good moisture conditions not only in the territory of the African continent, but also in other monsoon regions of Arabia, India, northern Australia and the equatorial belt of southern America (Padu2 and Quintao, 1982). Since with global warmings the temperature difference between the equator and poles decreases, westerlies also decrease in middle latitudes, which results in reduced water vapour fluxes from the ocean to the interior continental regions. As the reconstruction of the precipitation pattern for the date Riss - Wurm Mikulino - 5e orthoclimathem showed (Fig. 8.6), with larger warmings (when the global temperature rises by 1.5 - 2.0"C as compared with the modern) moisture conditions improve in the whole extratropical zone of the Northern Hemisphere. Probably, at higher temperatures a decrease in westerlies is compensated for by increasing meridional circulation, if the northern seas are warm (Drozdov, 1983), and increasing convective rainfall.

Resume ( I ) Analysis of proxy geological data (paleobotanical, oxygen-isotope, paleohydrological etc.) allowed us to reveal in the climate history of the last 18 ka about 30 nannoclimathems occurring synchronously and most strongly expressed in high latitudes. (2) The climatic history of the Late Glacial - Holocene can be divided into three large intervals: Anathermal described by a positive trend against which short-term (from 50 t o 200 years) kryochrons are recorded with a 2 - 3°C global temperature decrease; Megathermal, the time with relatively warm and stable climate; and Katathermal characterized by global temperature lowering trend and 'increasing climatic instability. (3) The complicated nature of the climatic changes for the last 18 ka is a result of the interrelation between many factors: solar radiation variations due to astronomical factors and atmospheric transparency fluctuations because of explosive volcanic activity, natural atmospheric carbon dioxide changes, alterations of underlying surface albedo due to continental and sea glaciation development, and varying areas of tropical forests as well as deserts and forests in middle latitudes. (4) Spatial surface air temperature reconstructions carried out for two time intervals - the glaciation maximum and most considerable Holocene warming, 6.2 - 5.3 ka BP - showed that the global temperature change in the glacial - interglacial cycle makes up 5 - 6"C, which is higher by an order of magnitude than the variations typical of the period of instrumental observations.

295

(5) Analysis of proxy data on moisture variations in different latitudinal zones during the Late Glacial - Holocene demonstrated that the global temperature - precipitation relationship revealed by data of instrumental observations for the last 100 years was also preserved in the past. During nannothermochrons the moisture conditions have improved in monsoon regions and north of 65"N, whereas southern regions in middle latitudes and northern regions in subtropical latitudes suffered from a rainfall deficit. With coolings the pattern was reversed. (6) The major mechanisms explaining the global temperature - precipitation relationship are associated with varying mean meridionai gradient and changes in the position and intensity of main atmospheric pressure centers.

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SUMMARY Global climatic events - an empirical basis for high-resolution stratification A comparison of the stratigraphic sequence of the continental Pleistocene of 12 Northern Hemisphere regions with that of oxygen-isotope stages (and substages) of the deep-sea Pleistocene has proved the possibility of identifying concurrent climatostratigraphic markers ( = signals, boundaries) in most of the sections and sedimentary facies. They are detectable by means of various methods in different assemblages and ecological facial environments. The use of a set of compatible and interchangeable correlation methods together with determining the age by physical techniques would allow their identification and detection over vast areas. The boundary levels are found to correspond to changes of temperature trends, coinciding with drastic changes in the global climatic system over relatively short time spans (from a few thousand to a few hundred years). The mentioned boundaries are known to cut many local climatosedimentary units such as glacial assemblages, till suites and even peat bogs. This allows the distinction of climathems (climatostratigraphic units) from their local counterparts called stratogenes and regional units which include well-known glacial and interglacial complexes and their corresponding “epochs” as defined by Jessen and Milthers (1928) and Markov et al. (1965). The boundaries of climathems are synchronous and those of stratogenes and glacial - interglacial complexes are asynchronous, that is they are transgressive in time, particularly longitudinally. Recently it has become possible to differentiate between global and regional climatic events, their corresponding stratigraphic units (like climathems and stratogenes) and the boundaries between them because high-resolution methods of regional correlations and datings have already been developed. The difference between global and regional climatic events has become a major premise in clirnatostratigraphy in its modern understanding. It is now quite obvious that unless climatostratigraphic validation is accomplished, and detailed climatochronology is developed, no studies of climate history including those of paleoclimatic reconstructions are trustworthy. The present book is the result of an empirical study. The concept of global climatic events has not been based on a-priori hypotheses but on the comparative analysis of a vast amount of direct evidence, particularly on the numerical age of climatic signals and the duration of climathems. These have been divided into three categories depending on the event duration. Thus kryomers and thermomers form nannoclimathems (NCT) spanning from a few hundred to a thousand years. Their sequence dated by I4C is proposed as the basis for the study of the Holocene climates, particularly those of the Late Glacial. Kryomers and thermomers with the time span of thousands of years corresponding to oxygen-isotope stages of Emiliani and Shackleton form orthoclirnatherns. Orthoclimathems (OCT) with the time span

298

of a few hundred thousand years roughly corresponding to the Alpine nomencluture (Riss, Mindel and the like) have been grouped into superclimathems (SCT). This approach t o the history of climate, based on a signaf indicating changes in the global temperature trend and the arrangement of climatic episodes according to their duration rather than t o the amplitude of temperature changes would provide the most universal basis for the comparison of different paleoclimatic data pertaining to various latitudes found on the land in the ocean and ice sheets. A timescale for the Northern Hemispheric climatic events is presented in Tables 6.1 and 8.3 and it is summarized in Fig. S. 1. This scheme has been updated and made more detailed than the earlier one (Zubakov and Borzenkova, 1983). It is obvious that no chronology is perfect, neither is it final. That is why it is proposed as another tentative scheme. That was the empirical part of the study. Many of the questions raised, however, still remain unsolved. One of them is the question of mechanisms and causes of rapid changes of the temperature trend, believed to be signals indicating changes in the global climatic system. We shall start our short hypothetical section by reminding the reader of some well-known concepts.

On the causes of climatic changes in the Pleistocene It is generally believed by climatologists and paleoclimatologists that the Cenozoic climatic trend is determined by two major factors, such as a decrease in the CO, atmospheric content due to weakening of explosive volcanic activity (Budyko et al., 1985), particularly after the middle Miocene, and a simultaneous increase in the Earth’s surface roughness, that is rising of mountains and deepening of the oceans (Monin and Shishkov, 1979; Frakes, 1979). Many authors believe that ocean deepening around Antarctica resulted in its permanent glaciation (Kennett, 1977) while uplifting of Tibet and the Himalayas led to the development of a distinct monsoon climate in the lower latitudes (Flohn, 1974, 1983). The major Pleistocene climatic variations manifested by the alternations of glaciations and interglacials are believed by many researchers to be attributed to the redistribution of solar insolation due to orbital variations (Milankovich, 1930; Yakovlev, 1956; Zeuner, 1959; Emiliani, 1967, 1978; Budyko and Vashishcheva, 1971; Zubakov, 1968; Sharaf, 1974; Berger, 1978; Fillon and Williams, 1983; Hays et al., 1976; Komintz et al., 1979; Kutzbach, 1983; and others). In fact, it was demonstrated in the book that the sequence of orthoclimathems is in a fairly good agreement with orbital rhythms of 40 and 100 ka, while that of superclimathems is in agreement with 370 - 425 ka rhythms. The Milankovich theory (1930) suggests (see Zeuner, 1959; Berger, 1978, 1979, 1981; Sharaf, 1974; Berger et al., 1984), though, that: (1) solar insolation variations due to variations of the Earth’s orbit are gradual and smooth, while the empirical data presented in our book give evidence of abrupt climatic changes; ( 2 ) solar insolation variations in the Southern and Northern Hemispheres are not

SCT

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Fig. S . I . Synoptic climatochronological scale for the Late Cenozoic. Age in ka. Temperature curve is the estimated temperature of the warmest month in the whole of the hemisphere). high latitudes (60&70"),Atlantic sector of the Northern Hemisphere (curve B 1 - normal polarity, 2 - excursions in the Brunhes orthomagnethem (see Table 2.5). 3 - reversed polarity. 4 - excursions: r l - Etruscan, r2 - Gotherburg, 5 - morphologically pronounced marginal glacial belts, 6 - reliable marine transgressions, 7 9 - datum planes: 7 first appearance datum, 8 last appearance datum, 9 - acme; 10 termination, beginning (a) and continuation (b) of the Late Glacial time, I I extinction of mammoth fauna, I 2 - Shuiner drift (Mesopotamia (1983). For the content of regional kryo- and therrnochrons (1 - 25 on curve A and nannochrons on curve B) see Tables 6.1 and 9.1. ~

%

300

simultaneous (due t o the e sin T effect) and frequently of opposite signs, while our empirical evidence indicates a complete synchronism of climatic episodes in both hemispheres; (3) variations of the Earth’s orbital parameters would include small temperature changes, not larger than tenths of a degree, while from empirical data the amplitude of summer temperature variations reaches 6 - 7°C in the glacial - interglacial cycle; (4) in the last 13 ka corresponding to a single cycle of growing and diminishing of solar insolation there should be no drastic climatic changes, while empirical data indicate global temperature variations with the amplitude of 5 - 6°C in a hundred years, e.g. between Late Dryas and Allerod, Allerod and Frisland, etc. (Flohn, 1974, 1979, 1985; Ruddiman, McIntyre, 1981 et al. etc.). Thus it appears that on the one hand climatic - sedimentation cycles are in agreement with the orbital variations as verified by statistical analysis (Hays et al., 1976; Komintz et al., 1979; Morley, Hays, 1981), the relationship being used for updating of oxygen isotope stages; on the other hand, the role of orbital factors in the changes of the Pleistocene climate remains “mystic” as Kerr (1984) figuratively wrote. This fact allows some authors - both geologists, (as the authors of the North American Stratigraphic Code, 1984) and climatologists - to reject the possibility of detecting climatic signals and thus depriving climatostratigraphy of its usefulness as a research tool. Many researchers attribute the changes of the Pleistocene climate to the changes in the atmospheric transparancy due to explosive volcanic activity and its resultant sulphur aerosols in the upper atmosphere. In particular, temporal correlations of short-term (1 - 3 years) coolings by 0.5 - 1 .O”C with strong volcanic eruptions have been established for the last century (Budyko, 1984; Kondratyev, 1985; Asaturov et al, 1986; Rampino and Self, 1984; Rampino, 1979 et al.). Other scientists do not accept this hypothesis. As indicated by the records shown in Fig. S.2 the correlation is also true for the Holocene to a certain degree. The datings of eruptions, though, as well as those of the Holocene kryomers are far from being perfect to suggest that volcanic eruptions would always precede coolings. It is not clear, however, in what way a short-term climatic impulse due to aerosol emission would induce a cooling that continued for hundreds of years, as in the middle Holocene, for instance (Fig. S.2). Let us assume that a large volcanic eruption similar to that of the Toba volcano, 100 or 75 ka BP, or a series of annual eruptions was one of the causes of the periodic prolonged Pleistocene coolings, though we do not have enough trusthworthy examples to prove it. The best evidence is found in the bottom cores of the eastern Mediterranean Sea which provide a regional inventory of ash layers dated relative to the oxygen-isotope stages (Cita et al., 1981). Most of the ash layers are found to correspond to kryorner portions of the section (Fig. 8.7). There is, however, abundant evidence showing thermorners to have volcanic signals as well. Thus the climatic transition from the Wurrn to the Holocene is marked by a number of ash layers with 14C ages of 9.8 ka BP found in the North Atlantic sections (Ia termination) (Fig. S.2). Their equivalent sulphate layers were found in the Greenland ice sheet section (Fig. S.2).

ka U2-

468-

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Fig. S.2. Summarized paleogeographical and palaeoclimatological characteristic of the Late Glacial and Postglacial. (a) Estimates of global fall-out of volcanic acids based on pH analyses on Crete ice core (Hammer et al., 1980; Hammer, 1984). (b) Paleotemperature record based on oxyden-isotope data on lacustrine carbonate from lake Tingstade Trask, Gothland (Morner, 1980). (c) July air temperature in the zone between 60"N and 70'" from pollen and other proxy data by the authors (1984). (d) Oxygen-isotope record from the North Atlantic (Ruddiman and Mclntyre, 1981b). (e) Caspian Sea level fluctuations (Rychagov, 1977; Svitoch et al., 1980; Varuschchenko et al., 1980; Zubakov, 1974). (f) Variations of the sea-level (Morner, 1973, 197911980; Fairbridge, 1977, 1983). (g) Holocene migrations of the warmwater front in north-western Pacific (Taira, 1979). (h) Recurrence of the high lake-level in the tropical zone of Africa by data of Street and Grove (1979). (i) Recurrence of the high lake-level of the Great Basin of the USA (Street and Grove, 1979). (k) A S - solar radiation (070 from present level, calculated by the authors using data of A. Berger ( 1 9 7 8 ~(I) C 0 2 - atmospheric CO: (ppmv) averaged by the authors using data from ice cores (W. Berger, 1983). (m) Distribution of the CaCO, content in the deep-sea cores from the western coast of North America (Balsam, 1981). (n) Positions of the terminations in the North Atlantic (Duplessy et al., 1981; Duplessy and Pujol, 1983).

2

302

The Pleistocene volcanic activity is assumed to be continuous, though it might be slightly stronger during kryochrons than in thermochrons, though there is no quantitative assessment of the difference. The only evidence that is available indicates that large volcanic eruptions coincide both with the most intense growth of ice sheets (that is with the end of thermochrons and beginning of kryochrons) and with their destruction. This is a preliminary conclusion that the authors came to in Chapters 8 and 9. It is in agreement with empirical data obtained at Kamchatka and in the North Pacific by scientists of the Institute of Volcanology of the USSR Academy of Sciences (Melekestsev, 1982). The volcanoes are known to be extremely sensitive to abrupt disturbances of the lithospheric gravitational equilibrium. That is why the growth and decay of ice sheets with their accompanying sea level changes amounting to hundreds of meters should theoretically involve stronger volcanic activity (Fairbridge, 1977; Rampino et al., 1979; Rampino and Self, 1984; Taira, 1983). The most important inference is that all well-marked boundaries of the Pleistocene should correspond to the onset of warming rather than of cooling. This was shown by Broecker and Van Donk (1970) who introduced the term termination for geologically rapid transitions from maximum cooling to maximum warming. It is these terminations that are major boundaries in the Pleistocene climatic history. The volcanic activity seems to have nothing t o do with these boundaries. Thus, summing up, we can claim that the relationship of volcanism and climatic changes is rather ambiguous. Obviously, volcanic eruptions would result in short-term coolings, which in turn would trigger more profound temperature decreases. Likewise drastic changes in the global energy distribution would induce volcanic activity. It is, however, clear that the volcanism above cannot regulate the climatic engine, though it is directly connected to it.

Fig. S.3. Percentage of carbon dioxide in the air bubbles in ice cores through the Antarctic and Greenland ice sheets. Present-day atmospheric CO, content is 0.033%. 1 - Byrd Station, 2 - Dome C and 3 - Dome 10, Antarctica (Lorius et al., 1979; Neftel et a]., 1982), 4 - Camp Century, Greenland (Berger, 1982; Neftel et al., 1982), 5 - averaged curve compiled by Borzenkova from the data of Berger.

303

Let us now consider some recent developments in paleocIimatoIogy related to the problem of natural short-term variations of atmospheric CO, and water vapour content which were not discussed in the authors' earlier work (Zubakov and Borzenkova, 1983). In 1980- 1982 two laboratories (Berne and Grenoble) developed procedures of CO, extraction from ice based on the dry extraction method. The CO, content ice samples of 1 cm3 was measured by means of highaccuracy laser spectroscopy. Sensational results have been published (Delmas et al., 1980; Neftel et al., 1982). The CO, content in the atmosphere during the Late Wurm kryochron was found to be 190-200 ppm. Some 15.0-9.0 ka BP the concentration of atmospheric CO, increased to 350- 400 ppm, and it fell to 270 - 290 ppm (Fig. S.3) by the end of the Holocene. Oeschger et al. (1985) made a detailed study of the CO, content in the ice core from the Dye-3 borehole in south-eastern Greenland. The Greenland section covering the last 90 ka of geological history was studied with a 30-year spacing. This high resolution of sampling allowed them to establish an uneven pattern of CO, changes with alternating stable levels. Fig. S.4 shows that at a three-meter ice core at a depth of 1897 - 1899 meters some 40 ka BP two modes with concentrations of 180 - 190 ppm and 260-270 ppm alternated with a main jump with an amplitude of 60- 30 ppm which is believed to occur only 100- 150 years. It is interesting that Dansgaard et al. (1984) also established this abrupt change in the 180/160 ratio curve. The shift of 6 I 8 0 reaches 5 - 6%0; this corresponds to a high-latitude air temperature change of a few degrees. This boundary seems to correlate well with those between Bugry kryomer and Kashin - Hengelo thermomer (see Table 8.3). A similar pattern of CO, and 180/160 ratio variations (Fig. S.4) is found at a depth of 1860- 1830 meters, roughly corresponding to the period from 37.0- 27.0

i897

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/a99

m Fig. S.4. CO, and 6 '*O values measured on ice samples from Dye 3, Greenland. (a) The results of single measurements of the CO, concentration of air extracted from ice samples (circles) in the 30 m increment corresponding to about 10 ka of accumulation, about 28 - 38 ka BP (see Fig. 2.11). Solid line connects the mean values for each depth. The lower curve shows the 6 I8O measurements done on 0.1. m core increments; (b) also in the 3 m layer represent an accumulation of about 1 ka, 43 ka BP (after Oeschger et al., 1983, 1985).

3 04

ka BP, that is the end of the Wurm thermochron. About five large changes of CO, concentration in the atmosphere with their corresponding temperature variations have been revealed in 10 ka. The CO, concentration changed from 180- 190 ppm during a glacial mode to 240 - 250 ppm in an interstadial mode, with an amplitude of 50 - 60 ppm. The duration of such changes ranges from 500 to 3000 yr. These fluctuations correspond to the changes revealed by Paterne et al. (Valladas et al., 1986) in the core KET 8004 (see Fig. 8.9). These two examples allow us to conclude that natural variations of atmospheric CO, amounting to 50 - 70 ppm in a hundred years occurred virtually concurrently with air temperature changes. This pattern of abrupt and rapid CO, fluctuations together with climate changes in the Wurm are found to be very similar to that in the Holocene and Late Glacial (Fig. S.2). As seen in Figs. S.3 and S.4, the atmospheric CO, variations in the last 13 ka from the "pre-industrial" level of 260 - 270 ppm typical of the end of the "Little Ice Age" to the present level of 345 pprn are only starting to exceed the natural atmospheric CO, variations recorded for the Holocene. Shackleton et al. (1983) and Shackleton and Pisias (1985) have shown that the differences in the isotopic composition of the carbon of plankton and benthic foraminifera tests, as was theoretically postulated by Broecker (1984), actually reflect the changes in the CO, concentrations in the surface waters of the World Ocean. Since micro-organisms prefer I2C (light isotope) to 13C (heavier isotope) the greater are the rates of carbon transport to the ocean bottom by the skeletons of dead organisms the higher is the 13C concentration in the surface waters and in planktonic organisms. As seen in Fig. S.5 the record of 13C concentration differences in planktonic and benthic foraminifera displays a marked correlation with the changes of CO, concentration in the ice cores. Shackleton and Pisias (1985) compared the curves of 6 I3C in the plankton and benthos with 6 I8O of the core from DSDP-504 (Fig. S.6) and with the solar radiation curve controlled by the orCO, ppm 150

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bital variations. They concluded that the CO, concentration changes in the surface waters, and in the atmosphere, precede 6 '*O variations and hence they precede the changes of the water and air temperatures, but it should be remembered that a11 these changes are preceded by orbital variations occurring some 1000 - 2000 years earlier. Thus if their conclusion is correct, it implies that the orbitally determined changes of the solar insolation should be a major trigger of the whole Pleistocene climatic system. The duration of the main Pleistocene climatic - sedimentation cycles is found to coincide with the eccentricity rhythms of 100 and 400 ka; the latter is believed to be the main forcing of the climatic engine. In fact, 100 and 400 year cycles are recognized in deep-water sediments of the Arctic Ocean and by a sequence of Antarctic fjord-terraces (Initial Geological Report of CESAR, 1985; Myagkov, 1979), that is in those areas where the cycles of the Earth's axis should be most pronounced according to Milankovich theory. Why is this so? There is no answer to the question until the mechanism of natural variations of the so-called greenhouse gases is understood. This first of all concerns CO, and H,O, their short-term (from hundreds to a few thousands of years) and long-term variations are associated with orbital changes. Short-term variations of CO, and H,O seem to be the main factors governing climatic changes; their manifestations in the stratigraphic section are defined as nannoclimathems. Longterm variations of CO, and H,O can be assumed to be the basis of orthoclimathems and superclimathem identification. It is only in this way that we would be able to explain empirically-established common features of these three units such as their global synchronism (global atmospheric redistribution of gases takes months and a few first years, while in the

l

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306

ocean a similar process would take hundreds and a few first thousand years), their distinct climatic boundaries (defined as terminations, steps, and jumps) corresponding to a rapid change of climatic environments, the strongest cooling giving way to the strongest warming. Apparently carbon cycles, changes of hydrogen isotopes and changes of the depth of carbon compensation, would reflect the temporal redistribution of CO, in the atmosphere/biosphere/ocean system. Together with the oxygen-isotope stages they represent the outcome of a continuous intergeospheric gas exchange. Thus, it becomes obvious that the ocean, a major CO, reservoir, is the key factor in the study of natural variations of atmospheric CO, and H 2 0 levels. There is no definite understanding so far as to what oceanic processes would control the atmospheric CO, content, though a majority of researchers believe them to be associated with the surface water biological productivity and the rate of vertical water circulation (Berger, 1982, 1985; Flohn, 1985; Sundquist & Broecker, 1985; Siegenthaler and Wenk, 1984; Sarmiento and Toggweiler, 1984). The larger is the productivity of surface waters and the stronger is the vertical water mixing, the larger will be the amount of hydrogen pumped from the ocean surface (and hence from the atmosphere) to the deep ocean. The productivity of photosynthesis zones in turn depends on the amount of nutrients (nitrates and phosphates) and the duration of the vegetative season. It is a well-known fact that the most productive areas in the World Ocean are those of relatively cold waters in high latitudes (e.g. Antarctic waters) and upwelling zones in low latitudes. These areas seem to control the CO, budget in the atmosphere. Weber and Flohn (1984), Siegenthaler and Wenk (1984) show that short-period CO, and H,O variations, with amplitudes ranging from 1 to 7 years should be associated with upwelling/downwelling events in the tropics (10"s - 4"N). In upwelling areas with water temperatures of 20 - 25°C the productivity of surface waters reaches 1000 g/cm2, while the dissolved CO, concentration is at its minimum. In El Qino years surface waters are unusually warm (28 - 29°C) and poor in nutrients, the productivity decreases 3 - 7 fold, while the CO, concentration dramatically increases by a few tens of ppm, part of the CO, being lost to the atmosphere. It is important to mention that the alteration of upwelling and downwelling causes changes of water vapour content in the low latitude atmosphere accompanied by changes in CO, concentration. Weber and Flohn (1984) found a direct relationship between the atmospheric H,O and sea surface temperatures. Decrease of evaporation rates in upwelling years in low-latitude areas would result in a precipitation decrease there. El R h o further contributes to evaporation increase and larger amounts of water vapour enter the atmosphere. Ramanathan showed that water vapour has a larger greenhouse effect than previously believed. Thus processes in the upwelling zone should be determined by variations of both CO, and H,O, the two greenhouse gases. The upwelling variations would be determined by wind changes which in turn depend on temperature contra'- between the poles and equator. Thus it can be inferred that upwelling variations are determined by air temperature fluctuations in high latitudes. Flohn and Weber (1984) believe that volcanic eruptions, by decreasing the

307

air temperature in high latitudes, trigger the whole system of natural variations of CO, and H,O levels in the atmosphere. This is an interesting conclusion, though a disputable one. We agree that the suggested mechanism would explain short-term variations of atmospheric CO, and H,O, rather than their medium-term (hundreds and first few thousand years) and long-term variations. This mechanism would be more closely related t o vertical and deep oceanic circulation rather than to wind patterns. Oeshger et al. (1989, Siegenthaler and Wenk (1984), Toggweiler and Sarmiento (Sundquist and Broecker, 1985, pp. 163 - 183), Duplessy and Ruddiman (1984) and Shackleton and Pisias (1985) state that high-latitude oceanic processes should be considered as a major cause of the variations of CO, and H,O content in the atmosphere. The above-mentioned researchers suggest various models and hypotheses for the processes. All of them agree on the necessity of further research. That is why the present authors would give only a schematic presentation of their hypothesis, aimed at better explanation of their empirical results presented in the book. Firstly it is important to better understand why major climatic signals (termination) by which we can distinguish between thermomers and their preceding kryomers, are globally synchronous and why they are so wellpronounced, even better than the lower boundaries of kryomers. Secondly, we would like to provide a better insight into the generics of such climatic environments as climatic optimum in the Holocene and preceding thermochrons (5e, 7, 9 and so on). Thirdly, we would like to better understand genetic differences of nannoclimathems, ortho- and superclimathems. Let us examine Figs. S.2 and S.3. They show a rise of the CO, concentration in the atmosphere to start 15 - 14 ka BP, i.e. 1 - 2 ka after the summer solar insolation starts growing in both hemispheres (Milankovich, 1930; Sharaf, 1974; Berger, 1978; Kutzbach, 1983). This CO, increase continued to 9 ka BP when the incoming summer solar radiation (As) reached its maximum, exceeding the present level by 7% (Kutzbach, 1983). According to the classification suggested by the authors, this period is called Anathermal. As shown in previous chapters, the lower boundary, 16- 15 ka BP, is marked everywhere by ice sheet melting. Ice sheets covered most of the Northern Hemisphere land masses. As shown by Hays et al. (1976), the extent of sea ice in the Southern Hemisphere during the Wiirm Ice Age was twice as large as it is at present. Most of the Atlantic Ocean was ice-covered, no bottom waters (NADW) were formed in the Norwegian Sea. Areas of strong diatomian and radiolarian ooze accumulation shifted a few hundred kilometers towards the equator. This made vegetative periods longer, and the ocean productivity was at its maximum, compared with that of interglacials. The concentration of aerosol particles of marine origin in ice cores from the Dome C hole in the Wurm layer is five times larger than at present at it is 20 times richer in particles of land origin (Petit, 1981). This, together with a sharp increase of eolian dust and desert sand particles found in low-latitude deep-sea cores (Sarnthein et al., 1984; Rea and Janicen, 1982; Romine and Moore, 1981), indicates a strong increase of wind circulation in the glacial time. At the same time, the rates of dense bottom water formation around Antarctica and their discharge via the

308

Vema Channel also increased (Ciesielski and Weaver, 1983). Stronger wind and ocean circulations would cause stronger upwellings. As shown by Siegenthaler and Wenk (1984), an upwelling increase from 15 Sv to 30 Sv would decrease the CO, atmospheric content by 25 ppm. Actually the strength of upwelling seems to have doubled in the glacial time. A decrease of the amount of CO, in the atmosphere in the glacial time would be accompanied by a similar decrease of the amount of H,O in the atmosphere. The climate would become arid not only in middle latitudes, where loess-like rocks appear, but also in the lower latitudes. This is indicated by an almost complete disappearance of tropical rain forests which survived only in a few refuges (Fig. 8.13). Thus, the temperature decreased gradually in the glacial time. A decrease of summer incoming radiation flux can be regarded as a trigger of a climatic change manifested by sea ice growth. A gradual increase of the rate of the atmospheric circulation and, which is more important, the rate of the oceanic circulation are major factors leading to stronger upwellings and an increase of ocean productivity which extracts CO, from the atmosphere. Thus the Pleistocene kryochrons would be an outcome of the joint effects of greenhouse gases (CO, and H,O), pumped from the atmosphere to the ocean, and the change of albedo by larger extent of sea ice and snow cover. What was the environment at the boundary between the Glacial and Anathermal, 16- 14 ka BP? An increase of summer insolation would trigger a shrinking of sea ice cover and melting of ice sheets at their southern edges in the Northern Hemisphere (Fillon and Williams, 1983). This brings about a certain rise of sea level, and fresh run-off would decrease surface water salinity, which as shown by Berger (1982, 1985) would slow the rates of vertical ocean circulation and formation of stable water stratification. Then, due to Worthington’s effect, “excessive CO,” would accumulate in the upper warmer and less saline waters of the ocean. The fresh run-off (including icebergs) is known to be irregular, occurring in a jump-like pattern. This is attributed to ice sheet surging, that is t o dramatic discharges of vast volumes of glacial ice into the ocean, raising of sea level by 1 - 2 meters, and probable by 5 meters during very short time intervals (tens of years and a few years?). This mechanism would explain short-term and medium-term variations of CO, and H,O concentrations in the atmosphere and their steplike trend during the Anathermal. Apparently the contribution of the sea ice albedo to the short-term temperature variations is very large. It is quite probably that surges would provoke volcanic eruptions. They would only complicate the whole pattern of climatic changes. The most important characteristics of the Anathermal (and the first quarter of all other orthothermochrons) is the different weakening of upwellings, caused by the slowing of the vertical ocean circulation. This results in the formation of stable warm-water layers (70 28 - 29°C) devoid of organic life in the equatorial area. The evaporation from the sea surface abruptly increases, and much more precipitation would fall on low-latitude land masses. This situation would predetermine the pumping of CO, and water vapour to the atmosphere and its related rapid warming of the climate.

309

What would happen at the boundary of the Anathermal and Metathermal, that is at termination (9 ka BP) Ia? As is seen in Fig. S.2, the summer insolation is at its maximaum (AS 7%). Summer temperatures almost throughout the whole of the Northern Hemisphere (except the North Atlantic sector, where there are Laurentian ice sheet remnants) would also reach their largest values; AT:, is almost 1°C higher then the present, AT, at 50- 55” would be 4- 3°C higher than the present values, while the amount of CO, in the atmosphere would begin to decrease. This situation can be explained by two factors, that is certain new changes in the ocean circulation and the appearance of a new reservoir for the atmospheric CO,, i.e. tropical rain forests. The complete disappearance of sea ice from the North Atlantic makes it into an area of bottom water formation (NABW). These North Atlantic bottom waters would be warmer (3°C higher) and more saline than Antarctic winter surface waters. Thus, on reaching the Antarctic they stimulate sinking of Antarctic cold surface waters, contributing to the vertical water mixing. The upwelling thus becomes stronger again in lower latitudes, evaporation decreases, the lower-latitude climate becomes more arid. This chain of events is well demonstrated by Fig. S.2 (see the frequency graph of lake levels). From this time (4- 3 ka BP) the grasslands of the Sahara start to turn into desert. Thus, it has been shown that all global climatostratigraphic boundaries can be regarded as signals indicating changes in the “operation” of the oceanic biological pump, responsible f o r the pumping of CO, and H 2 0 from the atmosphere to the ocean and then back to the atmosphere. In particular, the most dramatic and distinct climatostratigraphic boundaries of the Pleistocene (terminations) correspond t o dramatic increases of abundance of carbon dioxide and water vapour in the atmosphere (50- 100 ppm for the first few hundred years). The better solar illumination of high latitudes induced by orbital factors may be considered as a certain trigger causing changes in the biological pump operation at the temporal level of terminations, that is ortho- and superclimathems. Short-period rhythms (nannoclimathem fluctuations) spanning tens of years and 2500 years are not associated with the high-latitude illumination fluctuations. They should have some other trigger not known as yet. These fluctuations might be caused by volcanic eruptions, variations in solar activity, planetary gravitational effects (lunar effects in particular) and most probably by the “parade of the planets”, which as shown by Fairbridge and Hillaire-Marcel (1977) would occur every 1130 years, and as further elaborated by Peterson and Shnitnikov (1957, 1968), Karlstrom (1961), Hevri and Karlstrom (1974) and Fairbridge (1976, 1977), would exert a certain influence on the vertical circulation in the ocean. Other climatic cycles with periods of 11 - 13, 22,45, 80, 170-200, 550- 600 and 1700-2500 years have been established by geological evidence. These seem to be attributed to the solar effects o n the Earth’s atmosphere. As is well known, solar activity effects on the climate and weather were argued and considered questionable, since no direct agreement between solar activity cycles and weather were found. Lately opinions have started to change. A national conference on the relationships of weather events with space processes, sponsored by the USSR Academy of Sciences and the State USSR Committee for

310

Hydrometeorology and Control of the Natural Environment, was held in 1985. As reported by Loginov and Fedorovich (1986) a number of trigger mechanisms involving solar effects on the lower atmosphere were discussed at the conference, they were: (1) variations of the solar constant which according to satellite data changed from 1364 to 1375 W/m2 during the period from 1978 to 1983; (2) changes of the ozone layer due to ultraviolet radiation variations; (3) cloud cover variations due to electromagnetic and corpuscular effects on the lower atmosphere condensation nuclei.

Obviously all these factors would affect both the water vapour content in the atmosphere and the productivity of organisms participating in the photosynthesis, that is on the carbon cycle. Thus we can infer that the earth’s climatic engine is controlled by trigger mechanisms i.e. by orbital and solar factors affecting the Earth’s climate in a direct way and indirectly via the biosphere, that is via the carbon cycle.

Resume The evidence discussed in the book allows the authors to make the following conclusions. (1) There is a common feature of all climatic oscillations of different importance and duration (short-, medium-, and long-period oscillations); that is all of them, independent of their scale, are but an alternation of relatively long intervals of stable climate with short abrupt and some times even dramatic changes of the climatic environment. This allows us to distinguish climatic steps of different scales (and some times even climatic crises) with duration from tens of years to the first few millions of years. The major ecological reconstructions in the biosphere seem to coincide with them. (2) The change of sign of the global air temperature trend in the Late Pleistocene and Holocene has been found to correlate well with rapid (hundreds of years) and abrupt fluctuations of the atmospheric carbon dioxide and water vapour with an amplitude of 50- 100 ppm. A hypothetical mechanism responsible for rapid fluctuations of greenhouse gas concentration might involve a biological pump, that is processes of organic carbon accumulation in the ocean. The pump is believed to draw carbon dioxide from the atmosphere to the ocean and vice versa. The rate of pumping is determined by the ocean productivity and by the rate of vertical water mixing in the ocean. (3) What mechanisms drive the biological pump and determine the regime of its operation is not clear so far. Though a certain regime exists for tens or hundreds of years, as for instance the orbitally determined distribution of solar insolation in different latitudinal zones, which in turn would control the extent of sea ice in high latitudes and displacements of the areas of major organisms participating in

31 I

photosynthesis, while short-period fluctuations of carbon dioxide and water vapour in the atmosphere spanning tens of years and a few first years should be caused by some other factors which are still obscure. These fluctuations might be governed by solar activity effects on the ocean productivity not properly studied yet. (4) Volcanic activity can be regarded both as a trigger and as a consequence of short-period climate variation, though we would be reluctant to assume it to be the major cause of the latter. (5) Since the carbon dioxide balance in the atmosphere - hydrosphere - biosphere is determined by a complex combination of solar and terrestrial factors, obviously there could nof be a complete correspondence between real climatic rhythms and external (orbital and solar) trigger cyclicity. That is why the history and chronology of the past climates should be based on empirical geological evidence. ( 6 ) Paleoclimatic studies should be based on a better understanding of the role played by carbon dioxide pumping which is regarded as a driving belt of the climatic engine on the one hand, and high-resolution methods of stratigraphic correlation of the Pleistocene deposits and their radioisotope dating on the other. There is the possibility now to subdivide [he past climates on the basis of global climatic events and clirnatherns, which are their equivalents in the sections. (7) A new concept of afmo-hydrospheric ( = exospheric) regimes could be used to typify paleoclimates. The term should be understood as the totality of processes occurring in the atmosphere, ocean, biosphere and upper lithosphere. The interaction of these processes would determine a long-period (hundreds of millions of years) status of the Earth’s climatic engine. Two atmo-hydrospheric regimes alternated during the last billion years, namely the psychrospheric and greenhouse therrnohaline regimes which constitute the reality of climatic history in the same way as the alternation of summer and winter seasons is the weather history. All paleoclimatic events of lesser significance would be oscillations of the regimes in question. (8) The Pleistocene has been closen as an example for which a classification of the shortest oscillations of paleoclimates has been established. The same oscillations have been found to have stratigraphic equivalents, these are nanno-, ortho- and superclimathems. These latter units represent kryomer and thermoiner divisions of climatic and sedimentary cycles which are recognized on the inter-regional basis in rhe secfions. Ortho- and superclimathems are the most distinct units; together they correspond to cycles in eccentricity with rhythms in 100 and 400 ka, which trigger them. Nannoclimathems cover intervals from the first few hundreds to thousand of years. No triggering mechanism to be related to nannoclimathems has been suggested yet. (9) A comparison of records from 12 continental regions with deep-sea data allowed the authors to propose a tentative chronological scale of nannoclimathems for the Holocene, and another one of ortho- and superclimathems for the Pleistocene. These scales are based on signal (or event) methodological approach lo geohistorical periodization. It forms a real foundation for summarizing all the available evidence and for compiling paleoclimatic maps of narrow temporal sections of the Pleistocene.

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PART I1

PRE-PLEISTOCENE CLIMATES Main steps of the Late Cenozoic glacial-psychrospheric regime standing

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PREFACE Initially, chapters covering the Miocene and Pliocene climates of the Russian Edition had not been translated, since the authors had exhausted the contract volume of the book, but evidence on the Neogene was in excess. Elsevier Science Publishers, after reading the initial manuscript, suggested that the authors should add a brief review of the Miocene and Pliocene climates thereby increasing the publication volume. This additional work on Part I1 took from June to December 1987. Part I1 has an independent Reference list which does not repeat those references given in the Reference list to Part I and the authors remind their readers not to forget this. Section 11.5 was written by 1.1. Borzenkova, while the rest of Part I1 was written by V.A. Zubakov. Chapter 10, Sections 11.1 and the Summary were translated by A.Ya. Minevich; Sections 11.2 and 11.3 were translated by R.V. Fursenko. Yu.B. Monakhova assisted in translating other sections and in preparing the manuscript. The authors are sincerely grateful to the above-mentioned persons as well as to O.S. Dyagileva, L.S. Bogdanova and M.L. Budnik for their help in the preparation of the manuscript and figures.

Leningrad, 12 December 1987

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INTRODUCTION

The climatic environments of the Pleistocene, as dramatic as they were, can be thought of as typical of the glacial-psychrospheric regime (a dynamic type), that is a specific state of the Earth’s exosphere determined by cosmic factors. The onset of the Late Cenozoic “galactic winter”, however, extends far beyond the Pleistocene. In fact, it is worthwhile to investigate how far beyond it goes. The investigations of the past 15 years have shown the amplitude of Cenozoic climatic changes to be much larger than it was thought earlier; the history of Cenozoic cooling has been shown t o consist of a sequence of steps. The author will attempt to reveal the major boundaries of the last glacial-psychrospheric regime. The analysis is only preliminary, though, which is understandable, for the interval under consideration spans a period of about 50 Ma, and the climatic evidence for this long interval is not adequately systematized. The review should start with a discussion of an important issue of methodology, that is: in what way should one proceed or what is the most rational approach to the retrieval of paleoclimatic evidence from the recently accumulated information of the oceans and land masses, this information being vast in volume and nonhomogeneous in its content, and the question is how one should systematize it. This discussion is worthwhile since the stratigraphy and chronology of the Cenozoic in the past 15 - 20 years changed dramatically and it might so happen that not all readers are fully aware of these changes.

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Chapter 10 PALEOCLIMATES OF THE PRE-PLIOCENE CENOZOIC

10.1. The state-of-the-art of stratigraphy, geochronology and historic subdivision of the Cenozoic Stratigraphic divisions of the Cenozoic have been known to be regional while the present-day correlations of events in various regions have turned out to be inadequately grounded and unreliable. It is not accidental that it is in the Cenozoic that the most archaic units of the geohistoric record like the Tertiary and Quaternary are still used. The only global subdivisions of the Cenozoic up to-date have been epochs ( = series); their boundaries, however, have been determined quite roughly. Recently the boundaries of the Miocene and Pliocene have been recognized in various countries as ranging from 5 to 12 Ma before present. It is not clear whether that is due to the limitations in the identification of stratigraphic boundaries, or to the uncertainties in the isotopic datings. Obviously, paleogeographic and paleoclimatic reconstructions made on the basis of such a timescale, as for instance Sinitsin’s (1980) reconstructions for the Cenozoic of northern Eurasia turn out to be quite uncertain. The stages of the Paleogene appear to be most unsatisfactory as shown by Krasheninnikov (1982). The Chronological Scale of the London Geological Society edited by Snelling (1985) differs favourably from others. Timescales for the Eocene, Oligocene and Neogene worked out by a group of authors headed by Berggren are presented in Figs. 10.1, 10.2 and 10.3. Analyzing these scales, one can readily see that scientists engaged in the study of the Pleistocene (Jenkins et al., 1985) and those dealing with the pre-Pleistocene Cenozoic (Berggren et al., 1985a; 1985b) refused to subdivide the Cenozoic into the Tertiary and Quaternary and, like the USSR Intergovernmental Stratigraphic Committee, they have agreed to divide it into the Paleogene and Neogene, the latter including three epochs (the Miocene, Pliocene and Pleistocene). This important decision conforms with the concept which has been advocated by the author more than once (Zubakov 1961, 1974, 1977, 1978b, 1980). The author has suggested that the Neogene (and Phanerozoic eonothem) should be crowned by the Pleistocene, while the recent sediments should be included into a new Technogean eonothem. This is in line with the ideas of Vernadsky (1987) and academician A.E. Fersman, his pupil, who were the first to realize the geohistoric and geochemical importance of the noosphere ( = technosphere) as the beginning of an absolutely new stage in the Earth’s history. This concept will allow us to define the stratigraphy on a strictly logical basis as a science of the stratisphere and past biospheres. Apparently the studies of the technosphere and the Earth’s future cannot be made in the context of stratigraphy, though an unusual definition of the Quaternary (Anthrhopogene) made by some authors as “an initial part of the period which continues in the future” seems to assume a certain futuristic potential. As can be seen from Figs. 10.1 to 10.3, the present timescale of the Cenozoic

3 20

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323

represents a system of parallel and fairly detailed and specialized scales, such as a magnetochronological scale, several zonal micro-biostratigraphic scales (plankton zones according to foraminifera, coccoliths, and radiolaria), and regional mammalian ages corresponding to a standard scale of epochs and ages of EuropeanMediterranian type area. All these specialized scales are closely related with certain events while numerical ages of the boundaries of their units are determined by isotopic methods; they can also be compared to the scale of linear magnetic anomalies. 33 anomalies of normal polarity are known to be distinguished in the 80 Ma time interval according to the Lamont Magnetic Polarity Chronological Scale (Heirtzler et al., 1968). These anomalies maintain their width ratios through all the profiles. Thanks to this, it is always possible to recognize a sequence of four to six polarity zones and to approximately date any reversal, on the basis of the anomaly width and assuming the spreading rate to be constant. The Lamont Magnetic Polarity Chronological Scale was repeatedly revised. Its latest version proposed by Berggren (1985a, 1985b) is shown in Figs. 10.1 to 10.3. This version identifies magnetic anomaly 5 (8.92- 10.42 Ma ago) with the magnetic polarity epoch 11, rather than with polarity epoch 9 as accepted by other researchers and the present author. Berggren et al. (1985b) have proposed common indices for chrons ( = orthomagnethems) of the anomalies of direct and older reversed polarities (see Fig. 10.3). A Magnetic Polarity Chronological Scale is a major basis of the dating of numerous biological markers, that is the first and the last appearance datum (FAD and LAD) of microfossils in deep water and marine continental sections as well as for the ages of microfaunal zones. The latter are predominantly associated with numerical indices from older zones towards younger ones and separately for the Paleogene and Neogene (Figs. 10.1 - 10.3). Thus, recent stratigraphy and chronology of the Cenozoic are based on the biomagnetochronological records with high resolution (almost an order of magnitude better) in comparison with the stages of the standard stratigraphic scale. The latter have lost their past significance as standards of geological time. However, they remain a common stratigraphic language, as stratigraphic sections of almost all the European Cenozoic stages (see a special column in Figs. 10.1 to 10.3) associated with micropaleontological scales, are known to be more accurately dated. And now we shall return to our task of a more rational systematization of the rapidly accumulating paleoclimatic evidence. Its sources include micropaleontological data and continuous isotopic records. However, as surprising as it may seem, the climatostratigraphic scale of the latest geochronological scale of the London Geological Society is applied for the Pleistocene only (Jenkins et al., 1985). Why is it so? One of the reasons, and it may be a major one, is the methodological principle of the Znternational Stratigraphic Guide (Hedberg, 1976) which determines the stratigraphic units depending on the method of their identification as litho-, bio-, magneto-, soil-, seismostratigraphic and chronostratigraphic. It is stated in the North American Stratigraphic Code that the identification of clirnatostratigraphic units seems to fail. Thus in worldwide general practice climatostratigraphic units are being identified while according to stratigraphic theory developed by Hedberg (1976) and his

3 24

followers they hardly exist. Clearly, this is a problem of philosophy used. A pragmatic philosopher would not easily accept climatostratigraphic units, since there is no purely climatic approach in stratigraphy. It will be possible to make a decision o n a past climate through the interpretation of numerous facts established by a set of various methods. The same can also be said of the relative geological time. That will be a certain interpretation of the evidence available. Hedberg, however, distinguishes chronostratigraphic units as a type, and the author cannot agree with him on that. Hedberg’s stratigraphic classification is not the best basis for the recent stratigraphy. We shall not go into a detailed discussion of this specific problem here. The author (Zubakov, 1968b, 1978, 1980) advocates an alternative approach, the one based o n the event (= time signal) principle for a stratigraphic classification, it has conventionally been used by European geologists. It needs further development to meet the requirements of present-day practice. A stratigraphic classification should not be torn away from the triad of stratigraphy, chronology, and geohistoric periodization, for a stratigrapher seeks to achieve, as his final objective, the

Table 10.1. Stratigraphical classification from the point of view of event conception. After Zubakov (1978, 1980). Standard scale

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325

development of a most rational systematization pattern of all temporal features of the Earth’s history. Since a description of the planet Earth today is impossible without distinguishing different geospheres, the study of the stratisphere is impossible without a description of the history of paleobiospheres, paleolithospheres, paleoatmospheres, paleooceans and paleomagnetospheres. Obviously, we need specialized geohistoric and event-stratigraphic classification, which will provide us with a wider approach to the Earth’s and a more objective synthesis of all the data available. Such an approach is now being developed for each of the five geospheres of the Earth (Table 10.1). As early as 1924 Stille developing a division of the history of Tectonosphere used a number of taxonomically related orogenic phases and cycles which could be now called tectonic magmatic cycles or tectonothem ( = synthem). A hierarchy of units (magnethem) is widely used in magnetostratigraphy. It consists of several taxa, e.g. Brunhes orthozone, Tikhvin superzone, Kiama hyperzone and SO on (Molostovsky, 1987). Global eustatic cycles (eusthem) or depositional sequences (Vail et al., 1977) with letter-number indices are identified on the basis of seismic evidence. Zones and sub-zones are also widely employed in biostratigraphy. Thus the use of climatostratigraphic units of different grades as a means for the systematization of paleoclimatic data as proposed by the author (Zubakov, 1978, 1980) is not anything new in the practices of the present-day stratigraphy. The largest taxon of a climatostratigraphic classification, a trendclimathem, was defined in Part I of this book, its corresponding climato-oceanic cycles alternating in the Earth’s history continue for tens and hundreds of million of years, the shortest of them - nanno-, ortho-, and superclimathem - will correspond to the events of the Pleistocene. To identify climatic episodes and their corresponding sediments ( = them) during the Cenozoic an additional two or three or more taxa might be used. A hyperclimathem, corresponding t o the half of the 1.2 Ma cycle of the eccentricity, will be explained in Chapters 10 and 1 1 in its relation to the Miocene and Pliocene. It is, however, somewhat preliminary to define taxa for the conjectural climatosedimental cycles of a few first tens of millions of years. Apparently we shall limit our efforts to a mere recognition of typical boundaries ( = steps) of climatic changes and empirically established cycles ( = rhythms) or events, as proposed by Keller (1983), dividing the history of Paleogene oceans into “events of Paleogene sea stratification” (Pss 1, Pss 2 etc.). Summing up, we can claim that the development of a specialized climatostratigraphic - climatochronological scale or division of the history of climate and atmospheres into periods will be the most rational approach to the systematization of paleoclimatic records.

10.2. The transition from the greenhouse - thermohaline regime to the glacial - psychrospheric one, 48 - 38 Ma As can be seen from Figs. 1.4 and 10.4 the cold bottom water accumulation started in the Danian, some 65 Ma ago, though surface waters in middle latitudes remained warm during most of the Paleocene. A general cooling of seawater world-

326

wide started 48 Ma ago during the Lutetian, though Keller (1983) times this event by the foram zone (P-14) as 44 - 43 Ma ago (Fig. 10.5). There is evidence indicating a mountain glaciation in Antarctica in the Early/Middle Eocene. However, if such a glaciation did exist, its extent should be limited. Otherwise it is difficult to comprehend the possibility of mammal marsupials, lizards and even crocodiles inhabiting Seymour Island, Antarctica (to the South of 60"s) in the Late Eocene. In Europe the oldest tropical rain forests with Normopolles of the Paleocene/Early Eocene are replaced by a paratropical flora in the end of the Middle Eocene, which can be classified into forma genera (Boytsova and Pokrovskaya, in: Grossgeim and Korobkov, 1975; Alexandrova et al., 1987). At the same time in northern Eurasia there are abundant subtropical evergreen broad-leaved forests

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Fig. 10.4. Temperature changes in surface and bottom ocean waters according to oxygen-isotope analysis data on plankton (above) and benthic (below) foraminiferas. Numbers near the symbols are the numbers of deep-sea DSDP sites and other (2 last marks) research vessels (After Savin, 1977).

327

("Poltava flora" of Krishtofovich, 1959) consisting of nut-trees, beeches, bay-trees, magnolias, palms and subtropical coniferous forests. Paludal forests consisting of cypresses, nyssa and alder were widespread, forming the basis for coal accumulation. The geographical distribution of these formations in Europe is shown in Figure 10.6. The southern Mediterranean and northern Africa were covered by savanna vegetation in the Paleocene/Early Eocene and thus we may conclude that an arid climate with seasonally variable precipitation prevailed. In Central Asia 48 - 50 Ma ago fossil unintathere mammal fauna is replaced by a more advanced tapiridian fauna including first artiodactyls and rodents (Devyatkin, 1981; Demberlane 1984) testifying to a certain dryness of the climate. High-latitude continental floras of the Northern Hemisphere reflect a somewhat

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328

Fig. 10.6. Climatic zonality of the Cenozoic Mediterranean (after Aleksandrova et al., 1987). A - Danian - Paleocene, 65 - 55 Ma; B - Eocene, 55 - 40 Ma, C - Oligocene, 35 - 25 Ma; D Miocene, 12- 10 Ma. Climate: 1 - 3 - tropical (1 - equatorial, humid, 2 - arid, 3 - evenly humid); 4 subtropical; 5 - moderate; 6 - paleolatitude position according to floristic data; 7 - paleolatitude ~

-

329

controversial pattern of the climate. As shown by many investigators (Axelrod, 1984; Baranova and Biske, 1979; Budantsev, 1984; Volkova et al., 1984; Wolfe, 1980; and many others) the northernmost findings of the fossil paratropical flora abundant in hydrophilous and thermophilous ferns are recognized as belonging to the Middle Eocene. Forests rich in Castanopsis, Aralia, Rhus, Magnoliaeae, Glyptostrobus and even in Ficus and Paimae SubGIires are established by spore-pollen spectra of the Middle Eocene Anjou Formation of New Siberian Islands, 76”N (Volkova et al., 1984; Baranova and Biske, 1979). Budantsev (1984) refers to Chemurnautskaya fossil flora of northern Kamchatka consisting of Palmae Sabai voronkovi, epiphytes Asplenium nidis, Laurophyllum, Myrtophyllum and other thermophilous elements to the Middle Eocene. The foresaid authors believe that the peak of warming in the Beringian falls on the boundary between the Middle and Upper Eocene, that is 45 - 44 Ma ago. Thus, we can speak about a heterogeneity of the local climatic trend as compared to the global Early Eocene maximum, 53 -52 Ma. This seems to be attributed to the history of the Turanian - West-Siberian Sea. According to new evidence, the largest transgression of that sea linking the Tethys with the Polar Ocean occurred at the boundary of the Middle and Upper Eocene during the Nurolian ( = Bartonian), 43 - 40 Ma ago (Shatsky, 1984). Sediments of that time from Turgay Strait and from areas further northward display communities of foraminifera with Textularia carinatiformis and Bolivinopsis spectabilis, radiolaria together with Heliodiscus hexastericus and Spongotrochus paciferus, diatoms, and dinoflagellates with Pixilla oligocaenica and Distephanus antiguus and others. It is important that these communities inhabiting the Tethys contain a number of tropical species (Kozlova and Strelnikova, 1984; Sharafutdinova, 1984; and others). The closure of the Turgay Strait in the Middle Priabonian, about 38 - 40 Ma ago, concurred with a change of the fossil flora and fauna in Siberia and north-eastern Asia. The substitution of moderately warm Turgay flora for the Poltava paratropic one seems to be a manifestation of climatic effects of the closure of the strait, though it is more probable that the substitution was mainly due to a global cooling occurring 44-43 Ma ago. The formation of heavy bottom waters in the Eocene was most probably caused by two mechanisms. The first one is Mediterranean type, typical of Tethys epicontinental basins, and the second mechanism is a polar one, characteristic of western Antarctica and the western Arctic Basin, where local areas of mountain glaciation form on land in coastal areas and in the polar night months seasonal sea ice cover forms. Before 50 Ma ago cold bottom waters seem to accumulate solely in the western areas of the Polar Ocean. After spreading of the Norwegian Sea they overflowed into the North Atlantic (Berggren, 1978, 1985a). position according to paleomagnetic data; 8 - tropical belt boundary; 9 - marine basins. Paleofloristic regional division of the Cenozoic Mediterranean (after Aleksandrova et al., 1987). A D - see (a); Paleofloristic kingdoms: I - Euroamerica “Normapolles”; I1 - Central-Atlantic “Proxapertites-Protuceae”; 2 - province boundaries; 3 - equatorial rain forests; 4 - savanna; 5 tropical forests; 6 - Mediterranean tropical forests; 7 - subtropical evergreen forests (Poltavian flora); 8 - moderate mixed broad-leaved deciduous forests (Turgay flora); 9 - moderate coniferous and broad-leaved forests; 10 - sclerophilic Mediterranean vegetation. ~

~

330

The polar-type bottom water formation in the Arctic Ocean would occur only if there existed an inflow of warm high-salinity waters mainly through the Turgay Strait and the West Siberian Sea. This water influx became stronger 45 Ma ago; a large stream of warm waters from Turanian and South Russian basins reached the east coasts of the Polar Ocean. This contributed to the formation of a specific climate there with smallest annual air temperature variations and high humidity, while variations of illumination were seasonal. Such a unique climate would form only if the CO, content in the atmosphere were high. Frequent climatic fluctuations, associated with variation of warm water influx into the North Polar Ocean and polar day/night conditions would strongly accelerate the evolution rates in this region of the Earth. 40- 38 Ma ago there is the largest variation in the climate’s history. It occurred at the P l 5 / P I , zone boundary, or at Pss 4 stage (Fig. 10.5). It is well documented by all geological methods. Thus in the ocean it is represented by a dramatically rapid bottom water temperature drop: 4” in 500 k a and according to Shackleton possibly in 100 ka (Savin et al., 1975; Savin, 1977; Kennett, 1977, 1982). Studies of fossil benthic microfauna from shelf sediments of the Tasman Sea indicate a similar amplitude of water temperature decrease from 17°C in the Late Eocene to 11.8”C at the beginning of the Oligocene (Burns and Nelson, 1981). The cooling is attributed to the appearance of a circumpolar current around Antarctica as a result of the Tasman Strait deepening and opening of the Drake Passage, a shallow strait as yet. Such thermal isolation of Antarctica would lead to the formation of glaciers which is confirmed by K/Ar dating of Marie Byrd Land pillow lavas (forming in the marine environment or under the ice) as 42 k 9 Ma (Le Masurier, 1970). There are no reliable data of the dimensions and volumes of the glaciation. Verbitsky and Kvasov (1982) suggested that the Eocene glaciation of Antarctica was of the same size as the present Antarctic ice sheet, though Drewry (1978), Kemp (1978), Mercer (1978) and Knox (1980) claimed that plant life still existed on the Antarctic coasts. The only undeniable fact is a drastic intensification of cold bottom water flow forming at the Antarctic coasts. This is established by breaks in deep sediments (Lisitsyn, 1980; Frakes, 1979; Keller, 1983). A drastic decrease of the depth of carbonate compensation 38 - 37 Ma ago (from 3.5 to 5 km) was an important event (Kennett, 1982). Climatic variations in the Northern Hemisphere are known to be spectacular as well, if we consider the very rapid (in terms of geology) substitution of Turgay flora for Poltava flora which took only 2 Ma. In Western Siberia this substitution is concurrent with the Eocene/Oligocene boundary 36 - 37 Ma ago (Shatsky, 1984; Zaklinskaya and Laukhin, 1978), the climate also changed from humid “climate of alkali soils” and brontotherian fauna to a moderately-humid “indricotherium savanna climate” with draughts (Lavrov and Panova, 1984). In north-eastern Asia the transition from paratropic Tastakh flora to the Oligocene Omoloy flora is marked by numerous sections and a stratigraphic well on Ayon Island (Belaya and Terekhova, 1982). In the Kavinsk Trough (Magadan District) the concurrent transitional Kavinsk fossil flora has been found, with dominating coniferous and small-leaved plants together with rare relict Castanea (Baranova et al., 1979). On northwestern Kamchatka the cooling is indicated by

33 I

Irgirninayam fossil flora among which virtually no thermophilic plants were found (Budantsev, 1984). Climatic changes were less noticeable in South Alaska. Though Wolfe (1980) and Axelrod (1984) noticed disappearance of many thermophilic plants from Ruvenien flora there. A particularly well-marked climatid boundary between the Eocene and Oligocene is recognized in the key section of Western Sakhalin Island (Krasilov et al., 1986). This boundary is located at the foot of the coal member of the Nizhneduy Formation and is recognized by the substitution of moderate small-leaved flora with Trochodendroides for subtropical broad-leaved flora. This event occurred in magnetochron C13-dated as 37.0- 35.5 Ma ago. As shown by Krasilov et al. (1986) this boundary was marked, in particular, by the appearance of a 400 ka sedimentation cycle in the section, which is characteristic of the Pleistocene, as noted above, or of the glacial regime. A drastic increase of the rates of cold bottom water formation (NABW and AABW) 40-37 Ma ago would indicate that the cryosphere was formed in high latitudes and the psychrosphere appeared in the deep ocean. Fig. 10.4 shows the bottom water temperature to drop to 5 - 8°C. Obviously, a zone of moderate climate would form at the same time. In the Southern Ocean this moderate climate zone is represented by a siliceous sedimentation belt marking a newly formed Antarctic Divergence, an upwelling zone, rich in nutrients. Climate zoning, i.e. an increase of the pole-to-equator contrasts in the atmosphere and of the vertical temperature gradient in the Ocean, would indicate intensification and acceleration of the atmospheric and Oceanic circulations. Thus the Eocene/Oligocene boundary is the first and possibly the largest step towards a glacier climate regime. That was the time when the present patterns of atmospheric and vertical oceanic circulations evolved: with the change of salinity stratification of ocean water into a thermal one, and it is not a casual coincidence that this boundary also marks a very important step in the evolution of the cenozoic organic life, primarily marine biota (Pomerol, 1981).

10.3. Psychrospheric climatic regime of the Oligocene/Early Miocene, 37 - 29 Ma The Buchardt oxygen-isotopic curve clearly displays an Oligocene cooling by marine molluscs from the North Sea (Buchardt, 1978). The sea surface temperature decreased by 12- 14°C in the North Sea area 38-26 Ma ago. This is a more dramatic temperature drop as compared with that shown by Ca/Sr analysis of freshwater molluscs from inland water bodies in the USSR (Yasamanov, 1985). As shown by Keller's deep-sea data (1983), the Oligocene climatic history emerges as having a much more complex pattern. As seen from Figs. 10.4 and 10.5, the bottomwater temperatures increased slightly in the middle of the Early Oligocene (about 32- 34 Ma, foram zone P20), while at the P20/P21 boundary (31.5 Ma ago) the temperatures decreased again. Terrestral flora and fauna confirm an Early - Middle Oligocene warming (Fig. 10.7). This is the time of a wide distribution of the Turgay tropical flora over the Eurasian land masses.

332

Fig. 10.7. Central Paratethys palynoflora and climate in the Oligocene and Miocene (after Aleksandrova et al., 1987). 1 - floristic elements correlation in the Central Paratethys: 1 - thermophilic “Poltavjan” floras; 2 - Arcto-Tertiary “Turgay” floras; 3 - intermediate floras. 11 - climatic parameters: 4 - mean annual moderate; 2 - subtemperatures (“C); 5 total annual precipitation (mm). 111 - type of climate: 1 tropical. ~

~

Monotonous swamp mixed forests with a relatively high percentage (almost 50%) of thermophilic species such as Fugaceae, Juglans, Ulmus, Ilex, Acer, Aralia, Zelcova, Nyssa and diverse Tuxodiceae extended northwards to the Polar Ocean (Belaya and Terekhova, 1982; Kartashova et al., 1985). E.K. Borisova using Klimanov’s method estimates the mean annual air temperatures in the Early Oligocene as 18 - 20°C in the Penzhina River watershed (63 - 64”N); the precipitation sum exceeds 1000 mm (Korotky and Pushkar, 1985, pp. 109- 116). Brontotheridae of the Early Oligocene in Kazakhstan and Mongolia together with abundant swampy and forest environments also indicate that semideserts of the present day had a climate of wet subtropical savanna (Devyatkin, 1981; Lavrov and Panova, 1984). In the Late Oligocene/Early Miocene prolonged Chattian - Aquitanian cooling developed, manifested by two wave modes on the oxygen-isotopic curves. The first cooling wave has a double pattern; it occurred 31 -28 Ma ago and it is recognized by the evidence from ice cores from DSDP 77B (Keller, 1983; Keigwin and Keller, 1984) and DSDP 366 (Grazzini and Lointier, 1980). It is separated from the second cooling wave (25 - 24 Ma) by a short, though strong, warming concurrent with zone P22, about 27 - 25 Ma ago (the Middle Chattian). Poor and Matthews (1984) claimed that the extent of the Antarctic ice sheet in Oligocene time was underestimated by the isotopic evidence (Douglas and Savin, 1971, 1973; Savin et al., 1975) and that the glacial “noise” should be taken into account starting from 30 Ma rather than from 15 Ma during the calibration of oxygen-isotopic data. Le Masurier (1970) using K/Ar datings of palagonite lavas established two phases of glacier growth in Marie Byrd Land in that time, i.e. 3 1.3 f 2 and 22.2 - 19.4 k 1.6 Ma. Leckie and Webb (1983) studying a well log at the Ross Ice Shelf margin found that glacial-marine sedimentation began 26 - 27 Ma ago there, reaching its

333

maximum about 20 - 18 Ma. Vail and Mitchum (1979) pointed out a drastic sea level decline around 29 Ma and two decreases about 24 and 22.5 Ma. Kerr (1984) noticed that Vail and Mitchum slightly overestimated the amplitude of sea level changes (up to 300 m). The existence of a deep Chattian regression exceeding 100 m as shown by Loutit and Kennett (1981) is recognized in different areas around the world. It is concomitant with an isotopic shift and a final shaping of the Antarctic Circumpolar Current in the Drake Passage area, some 25 - 22 Ma ago (Cieselski et al., 1982; Kennett, 1982). All the cited estimates of Oligocene climate variations lack accuracy as yet. We can suggest, though, that the Oligocene Antarctic ice sheet (29- 28 Ma) was of the same extent as the present one; the Oligocene ocean circulation acquires features that are typical of the present-day ocean (Nikolaev, 1986). The evidence relevant to the Northern Hemisphere continents reveals a similar pattern. Authors studying vegetation composition of the Oligocene/Miocene boundary (about 25 - 24 Ma) point out a disappearance of thermophilic species (Fig. 10.7). On the Chukotka Peninsula, as shown by Baranova and Biske (1979), forests of subtropical type are replaced by deciduous plants with dominating birches (60%) alders, pines. Over the Omolon River watershed (the right-hand tributary of the Kolyma, 67"N) the vegetation is represented by small-leaved hard-wood flora, similar in its composition to the northern dark coniferous taiga (Kartashova et al., 1985). As shown by the data from the stratigraphic well on Ayon Island (80"N) a diatom complex appears in the Late Oligocene Strata with abundant cryptophytes such as Tetrocyclus ellipticus, Eunotia faba and others. This suggests seasonal sea ice in the Polar Ocean some 25 Ma ago (Kartashova et al., 1985). Further to the south in Western Siberia there is a swampy coniferous/broadleaved forest with Castanea, Fagus, Carpinus, Platanus, Juglans, Liquidambar, Zelcova, Nyssa and archaic Alnus and Betula during the Turtass (Chattian - Early Acquitanian) along the shores of Zhuravskoye Lake (freshwater sea). This vegetation is very rich in diverse Taxodiaceae and lianas. Nikitin (1984) discriminates several floristic levels in the Turtass - Abrosimovo Age. The earlier level (Lyamin) has a mean annual air temperature of + 15°C and an annual precipitation sum of 1000 mm, and the later level (Vasyugan-Early Aquitanian) has mean annual air temperatures ranging from 10 to 12°C and annual precipitation sums of 800- 1000 mm. The substitution of the indricotheriae - tsaganomys fauna for the brontotheriae, accurately K/Ar dated in Mongolia as 32- 31 Ma (the lower part) and 24 -21 Ma (the upper part) (Devyatkin, 19811, suggests a cooler and dryer environment of expanded savannas in the Late Oligocene in Central Asia and Asia Minor. Diverse fast-running odd-toed animals (perissodactyles): rhinoceros and digging rodents Tsaganomys, Tataromys and others were found to be abundant there (Devyatkin, 1981; Nesmeyanov, 1977). Thus, paleofloristic and -faunistic evidence indicate a seasonal climate in the middle-latitude Asia with hot summers and cool winters with seasonal unstable snow cover. Rayushkina (1979) studied in detail fossil floras of the Indricotherium horizon of Kazakhstan sections (49- 50"N). Her results show the Mugodzhar Plateau (the present-day semi-desert) to have been covered by subtropical laurel forest with abundant Taxodium and evergreens including Cinnamomum cinnamomuni in the

334

Late Oligocene. Present-day Kolkhida forests can be regarded as their equivalents. Rayushkina claims the winter temperatures to be about 2°C and annual precipitation sum ranging from 1000 to 2000 mm. These estimates are in fairly good agreement with Nikitin's (1984) record for Western Siberia. Sinitsyn (1980) believes that the winter temperature in Mugodzhar was 12°C with an annual precipitation sum of 500 mm. The Late Oligocene flora of Bukhtarma (Western Altai, 50"N) is found to be younger than the Mugodzhar flora. It is represented by deciduous and dark coniferous forests containing Zelcova, Juglans, Pterocarya and the like: Rayushkina maintains that the climate in the foothills near the Altai was humid with relatively cold winters, with heavy snowfalls and hot summers, the temperature being about - 15°C in winter and about 30°C in summer, and the annual precipitation amount being not less than 1000 mm. Sinitsyn's paleoclimate charts show T , as 8°C' T, as 25°C and precipitation as 800 mm (Sinitsyn, 1980, figs. 14, 15, 22). The difference in the estimates of winter temperatures and precipitation amounts is rather large to be attributed to either different reconstruction techniques or to the nature of temporal correlations. Sinitsyn wrote: "It is quite possible that cooling was not a simple and progressive process, but it combined periodically repeated cooling and warming phases with a dominating cooling trend. At present these phases have not been properly established yet, though they will be in future, when our methods are more refined" (1980, p. 87). This is fully confirmed now. Thus Nesmeyanov (1977) in the same Indricotherium horizon has already distinguished sedimentary/climatic cycles testifying to substantial displacements of landscape/climatic zones in the Late Oligocene. In the Late Chatian and in the Aquitanian (25 - 22 Ma) the Northern Hemispheric fauna was of Agenian - Agyspe type which corresponded to Mein's teriochron MNl - MN2 (Mein, 1975). Zhegallo (1985) described this low-diversity population oriented to poor plant communities, similar to the present-day light forest, shrubs and dry steppes. That was the time of the first appearance of Bovidae and Ochotona with remaining Zndricotheriae. Thus, on the basis of the paleofaunistic evidence, the climate on the Oligocene/Miocene boundary was fairly cool and dry with distinct seasonality. As shown by Harris (Program and Abstracts, INQUA, 1987, p. 182) there is an oldest moraine of the Northern Hemisphere in the St. Ellias Range, Alaska, K/Ar-dated as 22 Ma. The extreme diversity of landscapes and biomas seem to be due to a small portion of grass - herb communities and well-developed aggradation plain facies, reflecting strong orogeny (Alpine orogeny) and global cooling. Nikolaev (1986) reconstructs sea surface temperature field for the 50"s to 50"N belt on the basis of oxygen-isotopic plankton foraminifera data for a layer of 300 m deep. He shows the Oligocene sea surface temperatures in the 50"s - 50"N belt to have been 3.5"C cooler than they are now. In the Southern Hemisphere this difference is 1.9"C and it is 5.4"C in the Northern Hemisphere; at a depth of 150 meters this difference amounts to 3.3"C. Obviously such low temperatures of the sea surface would be one of the causes of the global aridization of the climate.

335

10.4. Early -Middle Miocene optimum, 21.0 - 15.3 Ma Hydraulic-piston coring allowed us to get continuous and undisturbed cores of loose ooze and to study sections in the areas of rapid sedimentation with temporal resolution of a few thousands of years. Oxygen-isotopic curves of the DSDP site 289 (Fig. 10.8) are good examples of such a high resolution. The work of Douglas et al. (1971, 1973), Savin et al. (1975), Blanc et al. (1980, 1983), Woodruff et al. (1981), Keller (1979, 1981, 1983), Barron and Keller (1982) and of others has greatly contributed to a better understanding of the paleoclimate of deep Miocene sediments. Haq (1981, 1982) subdivides it into eight acme-events or climatic semicycles distinguished by fossil microfauna (Fig. 10.9). Fig. 10.10 illustrates the author's interpretation of the state-of-the-art of Miocene geochronology. The Lamont Scale of linear magnetic anomalies (column 1) is according to Berggren et al. (1985), the indices are those proposed by Harland et al. (1982). The ages of reversions (column 2 ) are estimated by new K-Ar records of McDougall et al. (1984) for Iceland (from 7.17 Ma to IS. 13 Ma, see Fig. 10.10); the Miocene

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Fig. 10.8. The changes in isotope composition o f benthic foraminiferas Cibicides shells from DSDP site 289, Pacific Ocean, equatorial zone, 154"E.A., reflecting climate and bottom water fluctuations in the Miocene (after Woodruff et al., 1981). Foram-zone N I I - the suggested time of ice shield shift to the sea level along the Eastern Antarctic perimeter, 15.3 13.5 Ma corresponds to HCT 25 -23. Zone N 13, 12.5 - 12.25 Ma corresponds to HCT 21 (the time of Wrangel Mountains glaciation in Alaska). ~

336

figures given below are from Berggren et al. (1985) increased by 0.6-0.8 Ma. Column 3 presents numbering of orthomagnethems (orthomagnetochrons) by three versions: (a) Theyer and Hammond (1974) and Ryan et al. (1974), (b) Berggren et al. (1985b) and Andreesku et al. (1987), and (c) Pevzner (1986) and Molostovsky (1986). Plankton biozones are numbered conventionally, foram zones (column 4) by Blow (1969) and nannoplankton zones (column 5) by Martini (1971). The ages of the boundaries have been recalculated in their comparison with the geomagnetic field reversions, and since the latter have uncertainties of about +. 5 - 8% as follows from their comparison with the ages of magnetic anomalies (Berggren et al., 1985; McDougall et al., 1984) the same errors are inherent in the age estimates of the biozones. A set of regional stratigraphic schemes includes three marine schemes and three continental ones. These are the Mediterranean scheme (column 6), compiled accorTIME IN MA -5-

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Fig. 10.9. Summary of nannofloral migrationary patterns in the North Atlantic Ocean through the Miocene. DSDP cores are represented by the appropriate site number, and sample levels in these cores are represented by lines. The Miocene record shows four climatic warming episodes as indicated by lowand mid-latitude assemblage incursions into higher latitudes and four cooling episodes are indicated by expansion of Coccolirhus pelugicus (high latitude) assemblage into lower latitudes (After Haq, 1982).

337

ding to Demarcq et al. (1983) and Berggren and Van Couvering (1974); the western and Central Paratethys (column 7) scheme according to Steininger and Papp (1979), Steininger et al. (1987), Vass et al. (1987) and Andreesku et al. (1987), and the Paratethys scheme (column 8) according to Nevesskaya et al. (1986, 1987), Pevzner (1986), Molostovsky (1986), Muratov and Nevesskaya (1986) and Ganzey (1984). The Miocene history is better represented in a continental section by teriochronological schemes. The Table includes several of them, namely the European scheme (column 9) according to Mein (1979, Berggren et al. (1985), Demarcq et al. (1983) and Steininger and Papp (1 979), Kretzoi (1987); the USSR scheme (column 10) according to Zhegallo (1978, 1985), Korotkevich (1981) and Pevzner (1986); and the North American scheme (column 11) according to Tedford (1981) and Berggren et al. (1985b). All these schemes have independent radiometric age estimates, obtained by K - Ar and fission-track analyses. There are certain interregional biological markers, primarily such as dispersal events, first appearance datum (FAD), last appearance datum (LAD), which will allow one to establish direct correlations between various specialized scales. The Miocene paleoclimatic evidence summarized by the author (deep-sea and continental records) is shown in column 12 by hyperclimathems: alternation of numbered paleoclimatic stages, that is kryochrons and thermochrons, 0.5 - 0.7 Ma long (0.6 Ma on average). Together (in a pair) they correspond to the 1.2 Ma eccentricity rhythm, discovered by Sharaf (1974) and Berger (1977). Column 13 denotes by a triangle old Antarctic moraines K - Ar-dated (Le Masurier, 1970; Stump et al., 1980), Patagonia moraines (Mercer and Sutter, 1982) and South Alaska ones (Denton and Armstrong, 1969; Armentrout et al., 1978). Floristic levels are shown (by a small leaf) in Iceland (Akhmetiev, 1980), in Japan and the USA (Wolfe and Tanai, 1980; Ananova, 1985) and in north-eastern USSR (Fotiyanova, 1984; Fradkina, 1983; Chelebaeva, 1978; and others). They reflect Miocene optima (thermal optima). On the right-hand side phytostages according t o Menke (1975) are given, warming and cooling (c), shown in column 14, recognized in deep-water sections of the North Pacific (Barron and Keller, 1982). Their ages were recalculated in accordance with the new boundaries of the magnetochronological scale of McDougall et al. (1984). AS seen from Fig. 10.11, the climatic waves of Barron and Keller are in fairly good agreement with either hyperclimathem or their parts ( = superclimathem), corresponding to half of a 400 ka rhythm of the eccentricity. This material allows one to assume the existence of long-term climatic rhythms, of 3.7 Ma and about I 1 Ma long. They can be regarded as a sequence of particularly well-pronounced warmings (thermal optima) which continued for about 0.6 Ma and 1.8 Ma, separated by periods of moderate climate, 3.0 Ma and 9.0 Ma long. Only the period of 21 - 15 Ma in the Late Cenozoic displays a significant warming trend with increased sea surface temperatures of 3 - 5°C near the Equator, (sst corresponds to the Eocene ones) and Antarctica bottom water temperatures increased by about 2-3°C (Fig. 10.4). The latter together with the evidence on the global Middle Miocene transgression (not less than 25 m) around 17 - 13 Ma ago (Vail and Mitchum, 1979; Loutit and Kennett, 1981) allows one to suggest that Antarctic glaciers at that time melted at least by one-third. Direct evidence from Antarctica,

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339

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1986; Ganzey, 1984; Pevzner, 1986) with the age revised; 9 - 11: teriochrons (mammalian ages); 9 - in Southern Europe (after Mein, 1975; Demarcq et al., 1983; Lopez-Martinez et al., 1987); 10 - in the USSR (after Zhegallo 1978, 1985; Korotkevich, 1981; and Pevzner 1986); 11 - in North America (after Tedford, 1981; Berggren et al., 1985b); 12 - parts of climatic cycle 3.7 Ma long; and hyperclimathems (half of the cycle 1.2 Ma long) and their approximate age; 13 - climatostratigraphic markers dated by the K/Ar method: tills (triangle) and floristic levels (leaf) in Iceland (Akhmetiev, 1980), in the USA (Ananova, 1985; after Wolfe, 1980; and others), in the north-western part of the USSR (after Fotiyanova and Serova, 1987), right - Europe floristical ages (after Menke, 1975); 14 - warming (W) and cooling (C) waves in a deep-sea section (after Keller and Barron, 1983) with the age revised.

340

though, is controversial. Thus Harwood et al. (1983) found diatoms in the transAntarctic Mountains moraines. This seems to indicate that ice-covered basins of East Antarctica, being hundreds of meters below sea level, several times became large marine bays in the Neogene. On the other hand, La Masurier (1970) and Stump et al. (1980) have dated pillow lavas as 22.2 f 1.6, ).4 f 1.5,

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342

19.2 - 17.98 and 16.2 - 15.4 Ma. All these numerical ages turn out to be included in the time intervals corresponding to cryo-HCTs. The oxygen-isotopic curves (Fig. 10.8) indicate the Miocene thermal optimum to consist of six plankton - foraminifera zones (from N5 to N10 inclusive), i.e. 21 - 15 Ma ago. This is confirmed by oxygen-isotopic curves of site DSDP216 in the Indian Ocean (Fig. 10.12). In the Mediterranean Sea foram zones 5 - 7 are found to correspond to the Burdigalian, hence it is dated as 22.7- 17.0 Ma, while foram zones 8- 10 correspond t o the Langhian (Demarcq et al., 1983; Berggren et al., 1985b). These levels in Eastern Paratethys are correlated with the Sakaraulian, Tarkhanian and Chokrakian, respectively (Nevesskaya et al., 1986). Thermophilic and stenohaline foraminiferous and mollusc fauna inhabited both basins at that time. Large Miogipsina foraminifera, Steginoporella corals and some ichthyoderms (Tripneustress planus, Scutella pauIensis and others) are found to appear at northwestern coasts of Tethys, in particular. These species are known to inhabit the Indian and Pacific Oceans at present. Mangroves appeared at that time as well (Demarcq et a]., 1983, 1987). The Lower Rhine Brown - Coal formation in the history of the terrageneous European flora corresponds to the Miocene climatic optimum. The lower beds of this complex include fossil plants of humid subtropics and tropical forests with palms, Mastixia, Symplicos, Engelhardtia and others. The palms during the Burdigalian extended northward, reaching the Baltic Sea coasts. The Upper Brown Coal formation corresponding t o the Upper Langhian does not contain fossil palms and other thermophilic evergreens; the forests of this period consist of taxodia, magnolia and laurels with a growing portion of broad-leaved taxa (Menke, 1975). This sequence of climatic events of the Early - Middle Miocene is well studied in Japan by now at Tozawan sections (Ikebe, 1973; Ikebe et al., 1977). They are characterized by fossil marine tropical fauna which includes Miogipsina foraminifera, molluscs (Vicarya and Turritella and others), corals, and also by remnants of abundant land plants and this is very important (Daijima). As shown by Tanai (Wolfe and Tanai, 1980; Ananova, 1986) this flora consists of broad-leaved forests mixed with evergreens, their portion (up to 44%) being the largest at this period of the whole of the Miocene. They are Comptonia, Cinnamornum, Machilus, Lithocarpus and others and conifers which are found at Taiwan and in Southern China now (Glyptostrobus, Taiwania, Gunninghamia etc.). Beds with Daijima flora with fauna of zone N8 seems to correlate with Hyperclimathem (HCT) 28, while the Sirakawa - Nisikurasawa (= foram zone N9 - 10) flora is very similar to that covering Kusu Island now; it correlates with HCT 26. Miocene desiduous floras are very well studied at the coasts of Cook Inlet, Alaska (60"N), where the climatic optimum of the Miocene is characterized in general by the Middle Seldovien flora, which is K - Ar dated as 15.9- 15.8 Ma and 16.8 Ma (HCT 26 - 28), containing up to 34% of broad-leaved plants and evergreens with Pterocarya, Ginkgo, Magnolia, Eucommia, Platanus, Celtis, Zelcova, Fagus, Juglans, Nyssa, Vitis, Glyptostrobus, Liquidambar and others (Wolfe, 1980; Fotiyanova, 1984). The present-day Hokkaido Island flora is equivalent to the South Alaska fossil floras. Summer temperatures on Hokkaido Island are about 20°C and the annual temperature is about 6-7°C (Wolfe and Tanai, 1980).

343

The Miocene climatic optimum floras of the Asiatic USSR are well studied, including the Bolotnino and Kizi floras of the coastal areas in the Far East (Ablaev, 1979), the Upper Duy of the Sakhalin (Krasilov et al., 1982) flora of “Hedgehog horizon” in northern Kamchatka (Chelebaeva, 1978; Fotiyanova and Serova, 1987; Fradkina, 1983) the “Beech” - Ildikilyakh horizon in Iana - Kolyma Lowland (Laukhin et al., 1983), and the Kireyevskoe Late Abrosimovka flora of West Siberia (Nikitin, 1984). All these fossil floras represent poor deciduous floras of Turgay type though with numerous subtropical evergreens extending to the Arctic Ocean coast with characteristic dominance of beeches, hornbeams and hazelnut groves. Thermophilic elements are present in the forests of West Siberia and Aldan, such as Comptonia, Broussonetia and Taxodium (Glyptostrobus) (Fradkina, 1983; Volkova et al., 1984). They are absent in Yana - Kolyma Lowland, though Fagus, Podocarpus, Nyssa and Zelcova are quite common (Laukhin et al., 1983) there. In Iceland the Miocene climatic optimum is recognized by Bothna lignites, 5C anomaly dated as 17.2- 17 Ma corresponding to HCT 28. A very warm climate is found to be typical of magnetochron boundaries 16/15, 15.6 Ma, that is HCT 26. Akhmetiev (1980) studied the Selaurdalur flora from Iceland which is represented by such thermophilic taxa as Fagus grandifolia, Glyptostrobus europeaus, Metasequoia occudentalis, Woodwardtia sp., Ginkgo, Criptomeria, Magnolia, Vitis and others. Akhmetiev believes that the climate in Iceland during the Bothna - Selaurdalur stage is reminiscent of the modern climate in the Northern Appalachian Mountains (T, 1 1 - 13”C, C P of 900- 1200 mm). The Middle Miocene is the time of the collision between the African and Eurasian plates. At the same time the connection between the Tethys and Indian Oceans was broken, the orogenic activity was strong, and the climate in the Mediterranean Sea area and in Central Asia became more continental. The terrigenous fossil faunistic record indicates a growth of temperature and an increase in aridity. In Europe it is represented by the Orleanian (MN zones 3, 4, 5), in Central and Minor Asia by the lower Oshin - Kushuk - AraIo-Turmey associations, in North America by the Hemingfordian association with prevailing savanna/forest anchitherium fauna of first mastodons and numerous rodents (Demarcq et al., 1983; Devyatkin, 1981; Tedford, 1981; Nesmeyanov, 1977). As described by Zhegallo (1985) this fauna is almost twice as diverse in species compared with the Late Oligocene one. It is characterized by a higher concentration of forest bioms and a respective decrease of open space. Boreal and tropical forests almost join (Afro-Eurasian exchange), while arid zone bioms remain in a few refuges in Mongolia and Africa. Mixed forests grew in Beringia; these forests were inhabited by bears, otters, raccoons, beavers and squirrels. The appearance of tropical bats (Hipposideridae) and false vampires (Megadermitidae) indicate that a warm and humid climate dominated over the Earth. Paleolandscape and paleoclimatic reconstructions of the Miocene thermal optimum have been made by many investigators, Soviet and non-Soviet. The following Soviet ones should be mentioned: Grossgeim et al. (1967), Sinitsyn (1980), Khain et a]., (1979), Yasamanov (1985), Zhegallo (1985) and Ananova (1986). Ananova developed phytogeographical zoning of northern Eurasia and North America for the boundary between the Lower and Middle Miocene. The scheme includes six

344

latitudinal-climatic zones and almost 20 provinces and subprovinces. There was no tundra in the Northern Hemisphere, at the Polar Ocean coast there were smallleaved swampy alder/birch forests with addition of conifers and a few broad-leaved taxa. Most of Europe and West Siberia were covered by mixed broad-leaved forests with beeches, oaks, chestnuts, magnolias, and swampy taxodians with swampy cypresses, and nyssa. Swamps of this type extended southward to Northern Kazakhstan. The Mediterranean Sea area was characterized by a hot and humid climate with winter temperatures higher than 15"C, as indicated by mangroves and Sabal Palms growing there. Only the modern deserts in Africa and Arabia seemed to be covered by dry savanna vegetation at that time.

10.5. Paleoclimates of the Middle - Late Miocene 15.3 - 7.8 Ma At least four large climatic stages can be distinguished in the Miocene from the basis of the Serravalian and foram zone N11; these are two coolings and two warmings (Figs. 10.9 and 10.10) divided by a period of 1.5 - 2.0 Ma, each stage including three sub-stages lasting 0.5 -0.7 Ma. The East Antarctica ice sheet was formed 1.8 Ma ago. This was a very important event, concurrent with foram zones N9 - 13 (Fig. 10.8). Ice sheet spreading northward to the Southern Ocean can be dated as 15.3 - 15.0 Ma ago (Kennett, 1982, Woodruff et al., 1981). Whether the Miocene (HCT 25 - 23) expansion of the high-latitude Antarctic ice sheet formed as early as the Cretacious is not clear as yet. Kennett (1982) believes that the ice sheet formation before the Middle Miocene was hampered by the deficit of atmospheric precipitation. As it was suggested by Schnitker, favourable conditions for glacier growth appeared only after warm North Atlantic water masses (NADW) penetrated to the coasts of Antarctica. This happened only after submerging of the Faeroes - Greenland Range in the Middle Miocene. This is quite possible though, as shown by evidence of other kind, some 12.5 Ma ago the Faeroes - Greenland Sill was a portion of a land mass or a shallow-water zone (Shor and Poore, 1981). Berger and Vinsent (1986) emphasize that 6I3C changes by 1 per thousand in DSDP 216 in the Indian Ocean (Fig. 10.12) preceeded by 2 Ma cooling (17.5 and 15.5 Ma, respectively). This event called the Monterey event (it was first described at Monterey formation section, South California) is unanimously attributed to upwellings, that is intensification of deep-water ocean circulation. A noticeable decrease of CO, concentration in the atmosphere 18.6- 18.4 Ma ago (HCT 31) is considered to be a result of that ocean circulation intensification. Berger and Vinsent consider the 613C event a precursor of cooling. It should be noticed, though, that the extension of the ice sheet towards the ocean 15.3 Ma ago is separated from the Monterey event 18.4 Ma ago by a 3 Ma long warming. We see neither a temporal nor a causal relationship between the Monterey event and the glaciation expansion. The period 15.3 - 13.0 Ma ago is characterized by climate fluctuations reminiscent of the glacial -interglacial events of the Pleistocene. The amplitude of the sea surface temperature variations near the equator reaches 2 - 3°C at that time, and the bottom water temperature variations reached 4 - 5"C, the duration of a cycle

345

ranging from 100 - 200 ka t o 1 - 2 ka (Fig. 10.8). This period was unique because it was at that time that high latitudes and bottom waters experienced dramatic cooling, while sea surface temperatures in low latitudes remained stable and even grew slightly (Fig. 10.4). Savin et al. (1975) associates the enhancement of the PoleEquator temperature contrasts with the formation of the Antarctic Circumpolar Current and thermal isolation of Antarctic latitudes. According 10 oxygen-isotopic data, the MiddleLate Miocene cooling was comprised of three waves dated as 15.3- 13.5 Ma (HCT 25-23, foram zone l l ) , 12.8- 12.2 Ma (HCT 21, foram zone 13) and 11.6-9.7 Ma (HCT 19- 17, foram zones 15 - 16). All three waves are marked by hiatuses and changes in sea-water stratification (Fig. 10.1 1). The effects of these cooling waves are found on the Northern Hemisphere land masses; these are mountain moraine groups. For instance in the lower part of the Jakataga glacial-marine series, Robinson Mountains, Alaska (Armentrout et al., 1978), there are tills at the boundary between zones N13 and N12 which enables us to date them as 12.5- 12.7 Ma (HCT 21), while the tills in the White River water shed (Wrangel Range) are K - Ar-dated as 9 - 10 Ma (Denton and Armstrong, 1969), and they seem to refer to HCT. Ciesielski and Weaver (1983) studied cores DSDP site 51 1 and 514 in the South Atlantic, and found that the West Antarctic ice sheet first appeared in the ninth paleomagnetic epoch about 10 - 9 Ma ago (also HCT 17). Since then this ice sheet, being highly unstable, is believed to control through its ice shelves bottom current velocities. Thus erosion hiatuses in the South Atlantic seem to be caused by the acceleration of West Antarctic glaciers. The time of formation of the West Antarctic ice sheet and of its related Humboldt Current is also well established by a rapid disappearance of thermophilic molluscs on the coast of Chile. In Europe during this time, Fischbach terrestrial flora grew; this seems to correspond to hyperclimathems 25 - 20. The main feature of this flora is the prevalence of broad-leaved forests with the minor role of subtropical taxa. Taxodiaceae, Betulaceae Fagaceae and Ulmaceae were the dominant groups at that time (Menke, 1975; Akhmetiev, 1980). The evidence of deciduous floras (Bryaunslaykur in Iceland and Medzhuda in the South Osetiya) indicates a transition from a Miocene thermal optimum to a Middle Miocene cooling. There are no tropical elements in the Late Karaganian flora of Medzhuda at all; this corresponds t o foram zone N11 , HCT 25, 14.7 - 15.3 Ma. As shown by Avakov (1979) in the foothills of the Caucasus the vegetation was comprised of Myrica ungeri, Quercus lochytis, Myrtus; higher in the mountains the forests consisted of such humid evergreens as Engelhardtia brongniarti, Magnolia dianae, Cinnamomum schenscheri, and even higher there were mesophytes with Carpinus, Ulmus, Tilia, Cornus. All these species can tolerate lower winter temperatures (down t o 10- 15°C). Avakov (1979) reconstructs the climate for the lower localities with Myrica where winter temperatures are about 15", summer temperatures are about 26", with mean annual temperatures of 19°C and annual precipitation sum equal t o 700 mm. Bryaunslaykur flora (as described by Akhmetiev (1980)) concurrent with the end of magnetochron 14 (14.2 Ma, HCT 24) consists of mixed dark conifers, pine and birebi trees, with a substantial addition of broadleaved plants and Taxodiaceae. The Middle/Late Miocene Antarctic ice sheet has the same age as the Sarmatian

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and Serravalian of the Black Sea coast (Fig. 10.10). The Sarmatian boundaries are dated by fission-track methods in the USSR as 13.7 and 9.75 Ma (Ganzey, 1984). Similar ages are suggested by other authors (Steininger and Papp, 1979; Kojumdgieva, 1983). The Serravalian is dated as 15.3- 11.5 Ma. These two periods are known to be associated with dramatic changes in landscapes and orogenies, while the Tethys area experienced evaporative processes, acceleration of land drying and cooling. This allows us to speak of the Serravalian/Middle Miocene (Demarcq, 1987) and of KhersoniadLate Miocene (Kojumdgieva, 1983) crises in the Tethys basin. According to the latest works of LopezMartinez et al. (1987) this Late Serravalian cooling (HCT 21 or it may be 19(?) was the most severe one in the Miocene. The migration of mammalia at the boundary between the Astaracian and Vallesian most vividly marked the rapid dramatic changes of the climate and of the landscapes in the Northern Hemisphere some 13 - 11.5 Ma ago (terriochrons MN 8 and MN 9), while in North America these events occurred at the boundary between the Barstovian and Clarendonian. The forest fauna was the first to migrate (HCT 22), these were Gomphotheriun.r, Zygolophodon, and Cormohipparion, which moved over the Bering and Faeroes - Greenland bridges. Later, during HCT 21, Canidae and Geomydae moved together with rodents. They, together with paleoarctic species, spread over large areas in Eurasia and Africa in a relatively short time period (Zhegallo, 1985). This Early Vallesian fauna is fairly accurately dated as 12.4 Ma by magnetic reversal 11/10 in Kalfa and Buzhor (Pevzner, 1986). The Vallesian Hipparion datum level (12.8 - 12.3 Ma) coincides with the lithological changes: gray humid facies (Nekrasov Group) gave way to variegated, often green, gypsum-bearing clays with carbonate nodules (Burlin Group). Apparently a unique hyperzonal landscape - climatic environment formed (tavrov and Panova, 1984; Martynov et al., 1987). Thus during 3- 4 Ma of the Astaracian, cooling together with increased aridity resulted in the degradation of OrIeanian forest phyto- and zoocenoses of high and middle latitudes and their replacement by steppes and savanna - steppe cenoses of the Vallesian. The Sarmatean (Homerian in North America) cooling and aridization is broken by a relatively short warming (HCT 22 and 20) coinciding with the lower parts of zones N12 and N14. This warming, called “The Second Miocene Optimum” (Fig. 10.10) is reflected in the lower beds of the Etolon formation, Kamchatka (Sinelnikova et al., 1979). The Lower Medvezhkina flora of Korf coal-bearing member in Kamchatka (Chelebaeva, 1978; Fotiyanova and Serova, 1987) in Korf Bay, north-eastern Kamchatka (60”N) will be best correlated with HCT 22. The fossil-flora-rich sediments overlay andezites K - Ar-dated as 16 - 15 Ma; these layers are cut by diorite intrusions, K - Ar-dated as 14 Ma. Lower Medvezhinsk flora contains broad-leaved and taxodian taxa; there are already no evergreens. Fradkina (1 983) refers the Mamontova Gora floristic level from a well-known Mamontova Gora section on the Aldan River to the same time interval. The Agnevskaya flora of Sakhalin Island containing Clyptostrobus and Fagus (Cape Markevich Beds) is tentatively referred to HCT 20, the Oldrich Station flora of Nevada was K - Ar dated as 10.7 - 11.2 Ma (HCT 18?). The latter includes 35 coniferous species and flowering plants, such thermophilic taxa as Comptonia, Platanus, Zelcoua (Bratseva et al., 1984). The Bessarabian horizon of the

347

Sarmatian dated by Ganzey by means of the fission-track method as 12.24 k 0.97 can be related to the Second Miocene Optimum. This horizon yields numerous species of the Middle-Vallesian mammalian fauna (the end of teriochron MN 9), found in Kalfa, Howenegg, Eni-Eskihisar, K - Ar dated as 12.5- 1 1 . 1 Ma (Zhegallo, 1978, 1985). The third wave of the Late Miocene cooling (HCT 19- 17) is correlated with zones N15 and N6 and "8 - "9 in deep-sea section, also with the Earlier Tortonian in the Tethys and with the lower part of the Pannonian and Khersonian in the Paratethys. This cooling wave has been dated as 11.6 - 9.7 Ma. The glaciation in the Northern Hemisphere is recognized by glacial-marine sedimentation in Yakutat Bay, Alaska. The glaciers descended from the Wrangel Range and St Ellias. The ages of the till were K - Ar-dated as 10 - 9 Ma (Denton and Armstrong, 1969). This coincides with the top of the Khersonian, determined by the fission-track method as 9.75 Ma old (Ganzey, 1984), and with a new invasion of the Hipparion fauna from America into Eurasia. This is believed to occur in the end of the Vallesian, zone MN 10 (Ulas p h a e ) , K - Ar-dated as 9.85 - 8.9 Ma. Thih .Goling can be recognized in the flora. In Europe it has a corresponding Bredstedt stage of the flora evolution recognized by Menke (1975) in the sediments of the Gram. In north-western Europe polydominant forests with the dominant broad-leaved taxa and with an addition of pines, firs, spruces and taxodia appeared during this time (HCT 19 and 17). The Holmatindur Tuff formation flora in Iceland as described by Mudie and Helgason (1983) belongs to the middle of magnethem 9, that is its age is approximately 9.85 Ma. Small-leaved and coniferous plants dominate the flora, though there are a few Taxodiaeceae and broad-leaved taxa. This cooling in North America, known as the Late Homerian, is represented by Table Mount flora (Wolfe, 1980) as well as by Hellensburg flora (Washington State), which are K/Ar-dated as 10.1 Ma. There are no subtropical elements in these two floras. Over the USSR area there are many places where fossil floras are found. This fact indicates several stages of the Late Miocene cooling. These are floras of the Upper Medvezhkina and Classical formations in Northern Kamchatka (Chelebaeva, 1979; Fradkina, 1983), Khapchan and Lower Gusin floras of the north-eastern area of the USSR (Baranova and Biske, 1979; Fradkina, 1983), the Late Tavolzhan flora of western Siberia (Nikitin, 1984) and others. All these floras are dominated by smallleaved taxa and conifers, with a very limited addition of broad-leaved species. However, taiga forest at that time did not yet exist. A double wave warming complaetes the climatic changes of the pre-Pleistocene Cenozoic (HCT 16- 14), which can be regarded as the Third Miocene optimum. It is recognized in the upper part of foram zone 16 and in nannoplankton zones NN 10 and NN 11 of the deep-water section. In the Black Sea region the warming is recognized in the Cape Panagiya - Cape Zhelezny Rog key section as the Middle Maeotian (Ananova et al., 1985). Pevzner et al. (1982) claim that the lower part of the Maeotian is concurrent with polarity epochs 7 and 8, (8.8 - 9.0 Ma). According to Berggren et al. (1987) it is the boundary between polarity epochs 10 and 11, i.e. 10.64 Ma. Andreesku et al. (1987) also date the Maeotian by polarity epochs 9 and 10. Thus the age of the third thermal optimum in the Mediterranean Miocene is not

348

dated definitely. This uncertainty is due to ambiguities concerning this interval in the magnetochronological scale. The author believes this warming to occur 9.7 - 7.8 Ma ago. This warming correlates with a warm-water Tortonian transgression, which seems to be of glacial-eustatic origin, extending to the Black Sea, where its effects are represented by the invasion of Mediterranian mollusc Paphia abichi, Mytitaster incrassatus and others. This is concurrent with the Middle Maeotian (Ilyina et al., 1976). The Late Miocene optimum is recognized in the terrigenous fossil floras and faunas. In north-western Europe it has corresponding Garding flora (Menke, 1975). In Iceland this last optimum is represented primarily by broad-leaved taxa (more than 20%) in the forest vegetation - Mokoudlsdalur phytochron (HCT 16) when the number of dark conifers decreased, the fossil plants display such thermophilic species as Pterocarya, Fagus, ex gr. syivafica, Acer ex. sect. Piatanoides and others. Akhmetiev (1980) describes the Icelandic climate of that time as having temperatures of 8 - 1O"C, with the precipitation norm of 800 - 1000 mm. The Flora Hredavatn (HCT 14) indicates a cooler climate. The northern coasts of the Black Sea were covered by savanna - steppe vegetation (according to Ananova, 1974; Shchekina, 1978), being in HCT 14 with hot summers and relatively mild winters, with episodic temperature drops down to -5 to - 10°C, the conditions are good for Acacia growing. Hoderz flora in the Dzin-Dza River Valley displays evergreen laurels and palms in the eastern Caucasus (Uznadze and Tsagareli, 1979). This third optimum by teriochronological scale correlates with the beginning of the Turolian, i.e. zone MN 11 and the beginning MN 12, with the age 9.6- 7.7 Ma. The analysis of paleoteriological evidence allowed Zhegallo (1985) to conclude that mammalian fauna during teriochron Pikermi had the most diverse taxa during the entire Cenozoic. The inhabitants of steppes and savannas played the dominant role; these were the hipparion, boveys, rodents, hamsters, hyenas and vivveras. There was an exchange via the Bering Strait bridge: it was the inhabitants of half-open bioms - Leporidae, Dipoides, Hipparion - that moved from North America; Rhinchotherium, Ochotona, Felis and others moved from Asia. This testifies to the expansion of savannas and valley forests, which are indicated by Gazella, Tragoceros, and Chilotherium. However, an extremely high productivity of hipparion biom does nor correlate with the climate aridity. As stated by Zhegallo (1985) Pikermian savannas of Europe were more similar t o the present-day Uganda landscapes, precipitation amounted t o 1000 mm in the north of this area and to 1500 mm in the south.

RCsumC (1) The Eocene Temporal optimum (and of the entire Cenozoic) was 52 - 53 Ma ago in the Late Ypresian (zone P 9). Paleoclimatic reconstructions for this interval seem to indicate a greenhouse - thermohaline regime with warm sea-waters, deep water temperatures being not lower than 10- 13°C. The entire system of sea- air circulation and the distribution of landscape - climatic zones over the Earth was entirely different. The high-latitude climate was frostless, and humid enough, the

349

amount of precipitation being not less than 700-900 mm in the modern arid landscapes. The carbon dioxide concentration in the air exceeded the present one by a factor of 5 - 10. (2) The last 50 million years display a steady temperature trend, complicated by numerous steps with coolings of various intensity alternating with warmings. The cooling peaks can be approximately dated as 51 -49, 44-43, 39-38, 36- 34, 31-28, 25-24, 22.5-21, 19.4-18.5, 15.3-13.5, 12.8-12.3, 11.6-11, 10.4 - 9.6, 9.0 and 7.8 - 7.2 Ma. Accordingly, warming peaks occurred approximately 49-45, 43-40, 38-36, 34-32, 28-26, 23, 21 - 19.8, 17.2- 15.3, 13.5- 12.9, 12.2- 11.6, 11.0- 10.4, 9.6-9.0 and 8.4 Ma ago. These sequences are not sufficiently accurate, but they display a certain climatic rhythm which lasts for 3.7 and 11 Ma. In the Miocene this rhythm (Fig. 10.10) is divided into six climatic steps, when warm and cold climates alternated after approximately 0.6 Ma. (3) Numerous biostratigraphic divisions, particularly teriochrons, coincide (or are similar) by their temperatures. For instance, the duration of the Mein teriochrons will normally be 0.6, 1.2 or 1.8 Ma. This allows us to hypothesize that in the Neogene the evolution of the climate and that of the organic life, and particularly of terrigenous biota, is affected by an eccentricity cycle lasting 1.2 Ma. The author distinguishes alternating cooling and warming stages, lasting about 0.6 Ma, calling them hyperclimathem and suggesting the latter as the main unit in a presently developing climatic periodicity of the Late Cenozoic. (4) It should be noted that the largest conventional stratigraphic boundaries of the Cenozoic coincide with major steps in a temperature trend. Thus the Eocene and Oligocene boundary, 36 Ma ago, corresponds to the time when the psychrosphere was developing, i.e. a qualitatively new state of the ocean, which would predetermine further evolution of the atmosphere and biosphere. The boundary of the Lower and Upper Oligocene (29 - 30 Ma) would be concurrent with the largest drop of the Cenozoic sea surface temperatures in the lower latitudes and its associated global aridization of the climate. The Oligocene/Miocene boundary (24 Ma) would be concurrent with the shaping of a strong Antarctic Circumpolar Current, which thermally isolated Antarctica and foreordained a new rise of sea surface temperatures in low latitudes and subsequent thermal Miocene optimum. (5) A major Neogene optimum corresponding to the Early Langian (zone N 8 and HCT 28) occurred 17.2- 16.5 Ma ago. It was characterized by a weakening of the ocean circulation (the smallest number of breaks), by sea surface temperature rise. It was accompanied by a more humid climate of the Earth in all latitudes, by considerable melting of Antarctic glaciers and by oceanic transgression. We can further hypothesize that warming of the Earth’s climate an increase of atmospheric moisture in the Langian was partly due to an increase of the fraction of the Tethys warm saline waters in the general production of heavy bottom waters, which would result in an increase of ocean stratification and of the carbon dioxide content in the atmosphere by 3 - 4 times as compared with the present-day levels. (6) This pattern of the paleoclimate over the Earth 17.2 - 16.5 Ma ago can hardly be accepted as a paleo analogue of the climate expected in the end of the 21st century with a predicted tripling of CO, concentrations, since the orographic (the mountains were two to three times less high than they are at present) and oceanographic

350

environments over the Earth at that time differed substantially from the present. (7) It is much more rational, however, from the point of view of predictions of future climates to reconstruct climatic environments of the third (the last) thermal optimum of the Miocene; these are hyperclimathems 16- 14, covering the period from 9.6 to 7.8 Ma and corresponding to orthomagnethems 8 and 7, zones N N 10 and 11, and MN 11 - the beginning of MN 12.

Chapter I 1 PALEOCLIMATES OF THE PLIOCENE 11.1. The Black Sea standard for the Pliocene

The South Russia Pliocene (the Eastern Paratethys and adjacent areas) is one of the best studied in the world. Many sections in the area are familiar to the author, who has studied Plio-Pleistocene stratigraphy in the Black Sea area, in the Caucasus, and near the Caspian Sea. The detailed climatostratigraphic division of the Pliocene has been made on the basis of the records from this area. The actual data used for this purpose are discussed below. The marine section of the Eastern Paratethys in the Pliocene is divided into four regional stages (Fig. 11.1) recognized by N.I. Andrusov and his colleagues by the change of hydrological conditions and fossil mollusc fauna (Andrusov, 1965; Eberzin, 1940; Zhizhchenko, 1952; Ilyina et al., 1976; Nevesskaya, 1986, 1987). The most important sections of the Black Sea Pliocene in Kerch-Taman type locality and in Western Georgia (Guryia) have been comprehensively studied in the last 15 years. The work included also a paleomagnetic division (Zubakov, 1974; Semenenko and Pevzner, 1979; Imnadze et al., 1982; Ananova et al., 1985; Grishanov et al., 1985). It has become possible to date a composite marine section of the Pliocene relative to a magnetostratigraphic scale (Figs. 11.2 and 11.3). The division of the continental Pliocene has been successfully accomplished in the northern Black Sea coastal area, where in the valleys of the Danube, Prut, Dniester, Dnieper, Don and other rivers there is a step-like complex of high river terraces. Their alluvium is characterized by freshwater mollusc and mammalian fauna (Chepalyga, 1967; Khubka, 1979; Bukatchuk et al., 1983; Aleksandrova et al., 1984). The distictly stratified sedimentary cover overlying the terrace alluvium and marine sediments is divided by Veklich and Sirenko (1976) into a reliably recognized succession of buried soils and loess - clay horizons. In the course of studies two competing stratigraphic schemes have been developed. One of them gives priority to the mammalian fauna (Bukatchuk et a]., 1983; Aleksandrova et al., 1984); the other is based on the data o n soil stratigraphy (Veklich, 1982; Sirenko and Turlo, 1986). Fig. I I . 1 summarises factual data reflecting the above-mentioned approaches to the division of the Black Sea Pliocene'. On the basis of the available paleoclimatic

'

Three holes were drilled during Leg 49 of the Gloniar Challenger cruise (379, 380, 381) in the Black Sea. The core analysis results are difficult to interpret, though, due to the lack of magnetostratigraphic divi5ion of the section and also to the extremely diverse estimations of the ages of sediments, so the author cannot dwell on it further here (Ross and Neprochnov, 1978; Stoffers et al., 1978; Hsu and Ciovanolli, 1979; Zhuze et al., 1980; Andreesku, 1987).

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records numbered superclimathems (SCT) have been recognized. They reflect changes of paleoclimatic environments in the region. As seen in the figure, most superclimathems by their extent correspond to either litho- or pedostratigraphic units (or their parts). In other words superclimathems are stratigraphic units, distinguished on the basis of the event principle. Each superclimathem seems to have its corresponding stratotype, though this is not necessary. The given numerical age ___-

c-

+I,

Iiyichevsk 1

167

Chokhvota'

2 -i

11

i i

-I

t v petenyrr Car rumanu

3,

Istriya

3.06

M plonicu

3.25

-4

1

i

Aidor loess

4.48

Stolnicheny alluvial c y c l e

5 agoid ak fauna:

_cy_

Goureny oilvvial cycle

Dic c f m e p r h i n r - H i p p . cr'o55um

IlI.umm----. C

Karboliyo beds

.?

2

__ 6.95 7.19

Fig. 11.1. Stratigraphy and climatic events sequence of the Pliocene Black Sea region. 1 - Magnetostratigraphical scale after Berggren et al. (1985) and McDougall et al. (1984) data. The author's corrections after Berggren (1987) 2 - Black Sea marine section (the author's version); 3 - Alluvial terraces of the Dniester basin and fresh-water molluscs fauna (after Aleksandrova et al., 1984; Bukatchuk et al., 1983; Veklich, 1982; Khubka, 1979; Chepalyga, 1967); 4 - mammalian fauna (after Aleksandrova et al., 1984; Alekseeva, 1978, 1982; Baigusheva, 1984; Topachevsky et al., 1986); 5 - Ukrainian superficial assemblage (after Veklich, 1982; Veklich and Sirenko, 1976; Sirenko and Turlo, 1986), the author's age version; 6 - superclimathems and hyperclimathems, their calculated age.

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estimates of superclimathems are only very rough. They are based on correlation of sections with paleomagnetic markers and on the interpretation with the use of isotopic data and orbital rhythms of 400 ka reflected in sediments. It should be noted here that the estimation of the temporal duration of the Pliocene superclimathems is in agreement -vith the available age of 7 Ma for the Pontian base. This age has been established by a number of independent estimations. Firstly, the estimation was made by a paleomagnetic method indicating a reversal between polarity epochs 7 and 6, coinciding with the Ingulets Kryomer by our data (Fig. 11.2). The age of the reversal is estimated by similar schemes as 6.6 Ma (Berggren et al., 1985), 6.77 Ma (Harland et al., 1982) and 7.17 Ma (McDougall et al., 1984); that is, on average it can be estimated as 7 Ma. Secondly, the fissiontrack method gives the age of the Maeotian top in the trans-Caucasian region as 7.14 f 0.58 and the Pontian base as 7.07 f 0.6 Ma (Ganzei, 1984). Thirdly, Minylitha convallis was found in the Maeotian top layers. This species is typical of nannoplankton zone NN 10 (Semenenko and Lyulieva, 1982). The upper temporal limit of this zone is assumed to be 8.2 Ma (Berggren et al., 1985); in any case it cannot be younger than 7 Ma. Fourthly, the top of the Evpatoriya limestones in the Ingulets horizon (the Crimea) yielded such findings as the camel Paracamelus typical of zone NN 12 of Mein. The first findings of the camel were made in KhirgisNur section, Mongolia, at the top of polarity zone 7 (Pevzner et al., 1982; Vangengeim et al., 1984). Thus, the Paracamelus dispersion datum is regarded as the best marker of the Pontian base. It is dated as 6.7 to 7.2 Ma, depending on the choice of paleomagnetic scheme. And finally, fifthy, the Pliocene composite stratigraphic section displays 30 sedimentary cycles extending from the base of the Kimmer i a n

Pon t ion

Maeotian

Staqe Horizon Depth, m

L l't hology 0

0

0.00. 0

0000 0

..

0

0.

I

I

. ..

Ostro coda Nunnoplankton Phytoplankton zone Phytoplonk ton (left) spore - PoNen (right)

Pollen composition

7 t h epoch

1

6 t h epoch

1st;

Gilbert

Polari t y ibiack -reverse t d ) zone

Fig. 11.2. A key section of the Lower and Middle Pliocene Black Sea region, located in the touthern Taman Peninsula between Panagia Cape and Kutrya ravine (After Ananova et al.. 1985).

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Pontian to the boundary of the Chauda (SCT 5; 995 ka). Assuming these cycles to be equal to half the orbital rhythm of the eccentricity, with its assumed mean duration of 370-380 ka (Zubakov, 1966- 1968), this would be 6.7 Ma for the base of SCT 35; with the rhythm duration of 425 ka (Sharaf, 1974) it would be 7.375 Ma; and with the rhythm duration of 413 ka it would be 7.19 Ma. The latter age is in a fairly good agreement with empirical data. The inferred age of the Pontian base, as indicated by a climatostratigraphic scale, would be 7.19 Ma2, while the age of reversal from polarity epoch 7 to 6 (the upper third of SCT 34) is estimated as 6.830 Ma, which is 60 ka longer than the known estimate of 6.77 Ma according to Harland et al. (1982). Let us briefly consider the succession of climatic events during the Black Sea Pliocene (Fig. 11. l ) and that of the corresponding superclimathems. The events in the Early Pliocene are referred to the Panagiya Cape key section, Taman Peninsula (Fig. 11.2) which was comprehensively studied by a group of scientists headed by the author from 1979 to 1982 (Ananova et al., 1985). Nannoplankton from this section (two collections consisting of 70 species), as well as that of the neighbouring one (Zhelezny Rog section) described by Semenenko and Pevzner (1 979) was studied by Lulieva, whose results were published (Semenenko and Lulieva, 1982) without her consulting the author and without their accurate reference. The study of our third collection, made by Ushakova in the Institute of Oceanology, Academy of Sciences of the USSR, turned out to be less informative than that of Lulieva (1982). The stratifications of the Panagiya and Zhelezny Rog sections located at different limbs of a synclinal fold are similar but not identical. Paleomagnetic analysis results differ in detail as well. That is why a general interpretation of extremely important results of nannoplankton analysis (Semenenko and Lulieva, 1982) is by no means an accurate and objective one. This leads to discrepancies in the correlations of the Black Sea Pliocene with the Mediterranean Sea section made by the author and Semenenko. Let us consider the climathems in the Panagiya Cape section (Fig. 11.2) and in its neighbouring Zhelezny Rog section (Semenenko and Pevzner, 1979; Semenenko and Lulieva, 1982). The end of the Maeotian on the Taman Peninsula is represented by diatomaceous shale, containing abundant Actinocyclus ehrenbergi, Rhaphoneis m aeotica, Thalassiosira m aeotica and others , acri t ar ch Micrhystridium regularipilatus, Cymatiosphaera reficulosa and others found in the Upper Tortonian and the Tripoli Formation in the Mediterranean zone (Ananova et al., 1985). As indicated by Lulieva, this list should include discoasters, a leading species for zone NN 10 (Semenenko, 1984). The Maeotian top contains regressive facies abundant in coastal plankton fauna with Braaderudosphaera bigelowi; Cyprideis torosa, etc., together with the Leptocyther ostracod complex (Ananova et al., 1985). The complete Pontean on the Taman Peninsula includes five lithostratigraphic units (Fig. 1 1.2) and seven superclimathems (Fig. 11.1); thermo-SCT 35 corresponds to transgressive Evpatoriya Beds with ostracods different from those in the Maeo-

’

Another estimate is known to exist yielding this age at 5.3 Ma: it is based on the unproved reference of the Pontian to the lower part of the Gilbert magnethem (Chepalyga, 1980, 1985; Molostovsky, 1986).

355

tian (Pontoniella, Caspiolla, and other pannonian species). These beds also yield forest pollen spectrum, with 45% of arboreal pollen represented by Zelkova and oak (Ananova et al., 1985). Kryo-SCT 34 corresponds to regressive Ingulets beds, characterized by quite a different arboreal spectrum, i.e. 70% coniferous (rich in spruces and firs). Shchekina (1979), Veklich (1982), Sirenko and Turlo (1986) claim that the coniferous forest with spruce, larch and subordinate birch, oak and elm covered the southern areas of the Ukraine in the Early Pontian. Thermo-SCT 33 corresponds by its extent to Odessa sub-horizon yielding rich mollusc fossil fauna (Paradacna abichi and others) and ostracods (Caspiolla labiata, C. pontica and others). The appearance of a new acritarch complex with Gonyaulax reticulatum and G. pannonicus characterizes SCT-33. The forest pollen complex with a majority of broad-leaved taxa, particularly of oak, gives evidence of a warming (Ananova et al., 1985). As indicated by Georgian records (Ramishvilli, 1969) during the Novorossiysk subage as a whole (SCT 35 - 33) the Georgian climate was close to a tropical one, as evidenced by a number of palm species (Nypa, Calamus, Phy tolephas), Magn oliaceae, A raliaceae, ferns : Mat onia, Cyath eae, Dickso nia and others. As shown by Ramishvili, the air temperatures in July were about 22 - 25", in January they dropped t o 8 - 10" and the precipitation amount was more than 1000 mm. Valley forests with Nyssa, Acacia, Araliaceae, Liquidambar grew over the northern Black Sea coasts in the Early Pontian (Negru, 1979). Kryo-SCT 32 is represented by the Portaferian Horizon with the mollusc fauna Congeria subromboidea, C. subcatinata, Dreissena stefanesckui and others, typical of the Pannonian; ostracods are represented by Cyprideis litoralis, Trachyleberis a f f . azerbaidjanica and others. The pollen is completely restricted by herbs. SCT-32 coincides with normal polarity event within magnetochron 6; this gives an independent confirmation of its age as 6.54-6.36 Ma. The Bosporian substage of the Pontian with characteristic fauna of Valenciennesia and a new ostracod complex with Caspiolla liventali, C. venusta and others corresponds to three superclimathems. The Lower (SCT 31) and Upper (SCT 29) Bosporian yield well-pronounced forest pollen spectra and more abundant acritarch complexes with Gonyaulax, than the Middle Bosporian (kryo-SCT 30). Faluns are referred to SCT 30 and SCT 29: Their tops are normal polarity and this gives the age of SCT 29 as 5.97-5.76 Ma, which is concordant with the paleomagnetic chronological scale. The Bosporian is divided into three partiads as has been well demonstrated by Chelidze (1974) and Taktakishvili (1984) in Georgia. The Lower (Urtiya) Horizon with the Dacian immigrants (Congeria rhomboidea rhomboidea and others) display sediments formed in water column with high salinity and higher temperature, the Middle Bosporian yielding Dreissena cf. iniquivalvis and Lirnnocardium, corresponds to the freshening phase, while the Upper Bosporian with Paradacna abichi exhibits another salinity increase. The Upper Bosporian ostracod fauna is found to be rich in Cypridadae (Imnadze and Karmishina, 1980) and marine foraminifera "Akchagylian" (Elphidiurn stellaturn) appear. The Ukranian Bosporian has been characterized by Shchekina (1979) and Sirenko and Turlo (1986) by means of the pollen complex of broad-leaved forests including thermophilic and hydrophilic Taxodium and Glyptostrobus. The pollen of the latter, as well as Tsuga, Rhus, Yuglans and other thermophilic plants, is recognized in the upper faluns

356

(SCT-29) and in Arshintsevo section, Kerch Peninsula (Sirenko and Turlo, 1982, fig. 2). The ferrestrial section of the Bosporian corresponds to a thick Znamenka red soil complex (Veklich, 1982). Kryo-SCT 28 in the Panagiya Cape marine section corresponds to a continental break, fixed by iron sands dividing the Pontian and Kimmerian. By earlier data (Andrusov, 1965), iron sands yielded remains of Mastodon borsoni. The scheme of Veklich (1984) shows SCT 28 as corresponding to the Belbek clay horizon. The latter in the key section near Lyubimovka Village (the Crimea) exhibits two reversions, as indicated by Veklich (1982); in our opinion, they are associated with the reversed polarity event within magnethem 5. Palynological analysis of the Belbek clays, Belbek River Mouth stratotype (Sirenko and Turlo, 1985, fig. 3) gives evidence of the prevalence of steppe landscapes with steppe vegetation and pine/oak forests in the valleys. Shchekina (1979) interprets the floral evidence of the Ukranian Late Pontian as “cold xerophyte savanna-like grassy landscapes” which changes into semi-deserts eastward of the Don, winter temperatures by Shchekina decreased down to 5 - 10°C below zero. The Belbek kryochron is concurrent with the change of Taurida (Pontian) fauna of mammals by Sagaidak fauna with archaic species of Promimomys though gazelles, mastodons, deers, antelopes, camels and rhinoceros prevailed (Adamenko et al., 1986). The Middle Pliocene of the Black Sea region (the Kimmerian) is less studied than the Upper and Lower Pliocene, particularly its continental sections, with its old alluvial plains which have no connection with the present valleys. These are the Karboliya and Kuchurgan beds in Moldavia (Lower Poratian) with Levantine mollusc fauna (khubka, 1979), Yergeni waterdivide sands between the Don, Volga, Manych (Rodzyanko, 1977, 1981, 1984; Krasnenkov et al., 1987). The marine Kimmerian is divided into three substages. The Lower, or Azov, Horizon is represented by thick sand beds and clays displaying poor mollusc fauna (Paradacna deformis) together with ostracods, mainly located in two or three ore-bearing interlayers. The latter are represented by non-oxidized tobacco-coloured siderite - leptochloric ores, formed by the red crust of weathering discharge to the deep-sea zone, due to washing out of Znamenka, Ivankovka and Lyubimovka red soils. These beds, as well as three ore-bearing layers of the Azov Horizon, correspond to superthermochrons 27, 25 and 23 (Fig. 11.1). As claimed by Sirenko and Turlo (1985), the Ivankovka soils (SCT 27) are typical ferrosiallatno-calcic red soils of humid subtropics. This is confirmed by the forest pollen, spectrum, demonstrating widespread broad-leaved vegetation over the Ukraine (with fir, chestnut, platane, Podocarpus, mulberry (Morus)and such subtropical species as Carya, Rhus, Pterocarya, Zelcova, Nyssa and Taxodium. The temperatures are believed t o have been as high as 23 - 28°C in June and not lower than 8 - 10” in the winter months, and the precipitation amounted to 800- 1000 mm, i.e. the climate over the Ukraine was very similar to that in the subtropics (humid and alternatively humid). During the Lyubimovka Age (SCT 25) the climate was more continental and moderate (Table 11.1). The kryomers of the Azov - Salgir (SCT 26) and Oskol (SCT 24) are represented by yellow - brown pulverized clays (with traces of loess) which are porous, with a vertical columnar structure, frequently salinized and having gypsum druses in the

357

terrestrial section. Pollen spectra yielded by the section do not allow us to restore steppe and forest - steppe landscapes with the prevalence of Chenopodiaceae (40- 46%) and Artemisia (16-2170). Arboreal plants were represented by abundant birches, alders, oaks and hornbeams. Winter temperatures dropped down to 10 - 15°C below zero then, while the precipitation amount decreased to 400 - 750 mm (Table 11.1). Kutrya beds 1, 2, 3 in the Panagiya marine section correspond to these kryochrons. They yield neither marine flora nor marine fauna; instead they display a high pollen content, with prevailing grass (34 - 5 1'To), mainly Chenopodiaceae. The arboreal pollen indicates the presence of birch, alder and willow, with the addition of freshwater plants Typha cJ angustifoh, T. cJ latifolia, Potamogeton sp., Nuphar sp., N. pumila and others (Ananova et al., 1985). Thus the Black Sea during kryochrons 28, 26 and 24 seems to have become an isolated freshwater basin. Nannoplankton with Discoaster quinquiramus (NN 11) was recognized in the Azov Horizon by Lulieva. This species is known to have disappeared from marine sections in the middle of magnethem 5, that is 5.7 Ma ago (Berggren et al., 1985). The upper part of the section displays Ceratolithus ex gr. acutus and C ex. gr. rugosus, typical of zones NN 12 and N N 13. On the basis of these findings and the normal polarity of the Azov Horizon, Semenenko and Lulieva (1982) place the Miocene/Pliocene boundary at the base of the Kamyshburun Horizon. We cannot agree with this inference. Though Semenenko and Lulieva in their article of 1982

Table 11.1. Quantitive climate indicators for the Pliocene Ukraine, obtained from pedological and palinological data after Sirenko and Turlo (1986); (a) the Northern Ukraine; (b) the Crimea. - --.--________~_ X P imm) SCT Climatochrons of the Ta""lldl T,",,,, T ,L l l i C l Ukrainian scale i"C) ("C) i"C) - - - _ _ _ _~ _ _ _ -28 Belbeck (bib) 2 -3* 1200 - 1500 27 Ivanovka (iv) (a) > 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. ~

~

362

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.

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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.

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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.

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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

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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.

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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

9 8

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02

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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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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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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). ~~~

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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

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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. ~

~

~

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

-4

T O

0

SCTh rn.yr.

summer

:-%,

I0

.. .

E (3

Menop

Likhvin

-

HCTh

Mammalian ages

8

4

Eem

Don

7 M. YR.

Chauda

1 2

T0.35 L

~

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

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SANTEHNIAN I Basal CALA6RIANs.l.)

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3

-Mammoth

GI1 RFRT -Coch;t;

4

-

Nunivok

4 I

KEY LEVELS

-

WARM

( D a t a mainly f r o m Krnnett & Walkins, 1914: portly rcinterpretcd) M A N GA PAN IAN

_____-__--__---I , WAIPIPIAN ------_-___-__-/-

\

\

\

-

Exp0nsion.d Northern Hemisphere glaciations

PIACENLIAN

N IY/2V ZANCLEAN 1LIND R I V E R ,FCTION

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-

Miocene-Pliocene boundary srrafotype, Capo Rossella,Sicily.

Evaporite

phase

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N E W Z E A L A N D STAGES A N D PALEOCLIMATE CURVE

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------5R 62

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-

Coiling chonge My ucosruem~i.

'Blue m o r l s "

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. Gr. conom(uzeu d a tum in Mediterranean TORTONIAN

Fig. 11.16. Correlation between Late Miocene - earliest Pleistocene paleomagnetic polarity scale, climate curve, planktonic foraminifera1 zonation and Mediterranean and New Zealand stages, mainly using paleoclimate changes (after Van Goersel and Troelstra, 198 1).

391

coastline (Blackwelder, 1981, Lietz and Schmincke, 1975; Meco and Stearns, 1981) - all this shows that the volume of the East Antarctic ice sheet considerably decreased in SCT 2 1. The climate of the HCT VIII was not uniform in regard to moisture conditions. On the whole, they were more humid than at present, which is, for instance, indicated by the savanna character of the Moldavian, Csarnotan and early Blancan fauna. Particularly humid conditions were in the terminal phase of the Brunssumian “C”. At this time in Europe at middle latitudes the Sequoia swamp-forests spread widely, leaving behind the lignite horizons (Menke, 1975; Meyer, 1981). This “Sequoia Climate” with cool foggy summers and mild winters reminding one of the contemporary climates of coastal Oregon, was probably associated with sharp temperature contrasts in the air masses which formed over the cold Norwegian - North Sea area and the North Atlantic isolated at that time (SCTs 20 - 18) from the above-mentioned seas. It might be thought also that during SCT 17 the North Atlantic was especially warm because of the rise up of the Peninsula of Panama some 3.6 to 3.2 Ma ago (Crouch and Poag, 1979) and the intensification of the Gulf Stream (Fig. 11.16). 11.4.3. Palaeoclirnates of the Late Pliocene (Villafranchian) 3.65

-

1.0 Ma

It is seen from Figures 11.1, 11.5, 11.8, 11.12 and 11.13 that the same sequence of climatic changes can be noticed in all regional sections, the time interval from 3.65 to 1.0 Ma. However, it was impossible to compare them with each other without radiological and palaeomagnetic data. These data allow one to estimate roughly the duration of SCT 13 and to identify the most probable synchronous reference levels. These are SCT 6 coinciding with the normal polarity event Cobb Mountain (1.1 Ma), SCT 9 corresponding to the Olduvai event, the 12th beginning with the Gauss - Matuyama reversion and SCT 16 coinciding with the Mammoth event. Superclimathems 16 - 12 form one increasing cooling (Fig. 11.15) wave which is especially well pronounced in the Northern Hemisphere. The coolings represented three global fluctuations and were accumpanied in the belt 35 - 45”N by the greatest transgression of the lakes of the Great Basin, the Caspian Sea and the Mongolian lakes in the Pliocene. SCT 16 is synchronous with the Mammoth reversed polarity event (3.07 -3.17 Ma). This was the time when the Greenland ice sheet and Iceland ice caps were building up (Akhmetjev, 1980; Berggren, 1978; Berggren and Van Couvering, 1974; Gladenkov, 1978) and when the Akchagylian transgression and the Great Basin pluvials were setting on (Smith, 1984). SCT 14 was the time of the first ice-sheet glaciation of North America (Easterbrook ad Boelstorff, 1981) and of the first advance of the Alpine (Cararo et al., 1975), Caucasian and Pamirs glaciers to their foothills (Milanovsky and Koronovsky, 1969; Zubakov, 1974; Nikonov, Pakhomov, 1984). During the 12th SCT the Laurentian ice sheet reached its greatest size (Easterbrook and Boellstroff, 1981). The Scandinavian ice sheet probably reached the southern coast of the Baltic at that time (Grichuk, 1985). The maximum of the Akchagylian transgression occurred just at the same interval (Zubakov, 1974). The three coolings indicated are clearly seen in the Black Sea basin

-

RANGE OF P E R T I N E N T

(elm. 4969, BnPR"" 8

R A N G E OF PERTINENT ATLANTIC SPECIES ( K U W S , 197

w W N

M A I N E V E N T S L E A D I N G TO E M E R G E W E OF ISTHMUS OF P A N A M A A N D F I N A L REDEPOSITION OF A M P H I S T E G I N A GIBBOSP. ON F L A N K OF S A N T A ROSA-CORTES RIDGE

S E Q U E N C E OF P L A N K T I C FAUNAS

IN CALIFORNIA STRATA (Atter Inch, 1967, 19731

Soto, 1976)

I S T H M l A N B A R R I E R COI'IPLETE; &, SUBSECUENT C O O L I N G K I L L S REDEPOSITED.

/

I S O L A T E D OFF C A L I F O R N I A : R E E F ; FAUNA ERODED ANII

E Q U A T O R I A L P L A N K T I C S SIJRGE NORTHWARD A G A I N ACCOMPANIED B Y A. GLBBllsB THEY M I G R A T L D ACROSS SHALLOW I S T H M I A N PASSAGE AND O C C U P I E D SANTA ROSA-CORTES RIDGE. B A S I N 20 F l l L I N G AND U P L I F T I L l M l N A T E A , GIBBQ5.4 FROM NORTHERN _. G U L F OF C A L I F O R N I A .

IJ/

~

25-APPROX

30,

35\

M A l E F I R S T I N D I C A T I O N OF T E C T O N I C U P L I F T AT DSDP S I T E ( C O L O M B I A B A S I N ) R E F L E C T I N G R I S E OF I S T H M I A N B A Q R I E R . I N C R E A S E D V E L O C I T Y OF G U L F STREAM D E P L E T E S S E D I M E N T S REACHING BLAKE PLATEAU.

144

E Q U A T O R I A L P L A N K T I C S SURGE NORTHWARD O F F C A L I F O R N I A , D E P O S I T I N CARYEN, I A P E R I A L , r\ND M I O - F E P N A N D O F O R M A T I O N S . A. M I G R A T E S FROM C A R I B B E A N TO GULF OF C A L I F O R N I A , I S T H M I A N PASSAGE CLOSED TO EXCHANGE OF D E E P AND MID-WATER TROPICAL PLANKTIC FORAMINIFERS. ENDEMIC PLANKTICS D E V E L O P I N A T L A N T I C - C A R I B B E A N AND DISAPPEARS THERE

~ ~ H A P P R O X I M A T ET I M E OF F l Q S T P L I O C E N E SUQGE O f E B U A T O R I A L P L A N K T I C F O R A M I N I F E R S NORTHWARD I N T O C A L I F O R N I A WATERS. NO E V I D E N C E O f A. THERE Y E T REPORTED.

Fig. I I . 17. Summary diagram of species ranges, paleomagnetic timescale, biostratigraphic zones and sequence of migratory and tectonic events related to emergence of the Isthmus of Panama (after Crouch and Poag, 1979).

393

in the horizons of loess-clays in subaerial formations (Veklich, 1982) and in the horizons of water-freshening in marine sections. Enormous ice-sheet glaciations and pluvial transgressions of the time interval 3.25 - 2.22 Ma were apparently the consequence of the isolation of the Polar basin from the Atlantic by the Faeroes - Greenland ridge. Actually the deep-sea drilling data of the Globar Challenger, Leg 49, showed that between 4 and 3 Ma the depth at the ridge did not exceed 100 - 200 m (Shor and Poore, 1981). Therefore it might be suggested that the Norwegian Sea and its gulf, the North Sea, were both turned into a freshwater basin covered with sea ice - a sort of “ice-water bottle” set against the European continent. A second factor was the temporary reestablishment of the Panama Strait which occurred between 3.2 and 2.5 Ma (Crouch and Poag, 1979) and this might have weakened the Gulf Stream. These two factors caused a shift of the cyclone tracks far to the south; meanwhile the sea-ice area in the Northern Hemisphere increased (Blanc et al., 1982; Rea and Schrader, 1985). Up to now it has not been possible in any way to determine exactly from radiometric ages the onset of the Calabrian transgression - the Selinuntian superstage which is usually associated with the penetration of “northern guests”: Arctica islandica, Hyalinea baltica, Cyteropteron testudo, Neogloboquadrina atlantica and others into the Mediterranean. The range of estimates varied from 1.64 Ma (Colalongo et al., 1981: Tauxe et al., 1983) to 2.0-2.3 Ma (Arrias et al., 1980; Nakagawa et al., 1980; Driever, 1984). This is the most uncertain time interval which could have been chosen for the Pliocene - Pleistocene boundary (!). We think that the appearance of the “northern guests” in the Mediterranean is explained not by climatic factors but by tectonic causes because of the gradual sinking of the Faeroes-Greenland ridge, which started 2 Ma ago or possibly about 2.3 Ma ago (Shor and Poor, 1981). This distribution of the “northern guests” continued also during the warm time intervals. This accounts for the extraordinary combination in the Mediterranean area of a subtropical climate on the continent with the cold Norwegian current which brought Hyalinea baltica to the coasts of Italy some 1.8 - 1.3 Ma ago, i.e. the 9th and 7th SCTs. Therefore, the discovery of the “northern guests” in the Mediterranean sections, which reflects the local tectonic situation in North Atlantic and the Mediterranean, cannot be an effective tool in making the global time correlation and determining the Pliocene - Pleistocene boundary. Also, it was during that period of time that the Gulf Stream penetrated deep in the Arctic Ocean and the Arctic basin was ice-free according to the unique data of Funder et al. (1985) on Cap Cobenhaven section, North Greenland (Fig. 11.18). Five superclimathems, from the 1lth to the 7th with ages of 2.2 to 1.2 Ma, are characterized by warm moderate climate as compared with the HCT V, which was however highly unstable. Frequent short-term climatic fluctuations have been revealed through pollen analysis of the sections of the Beringian (Fig. 11.19), the Caucasus (Shatilova, 1967, 1974, 1984), the Netherlands and the Central French Massif (Brun, 1973), West Germany (Menke and Behre, 1973), the Ural foothills (Nemkova and Yakhimovich, 1976) and Byelorussia (Makhnach et al., 1979, Levkov et al., 1987). Advances of mountain - valley glaciers from the Patagonian Andes (Fig. 11.13) were also quite frequent (Mercer, 1976). However, during this time, including the Eburonian, there were no large ice sheets. The warm climate of

394

HCT IV was probably the result of Gulf Stream inflow into the Polar basin, especially its eastern part included. This is confirmed by the appearance of the North-Atlantic molluscs Portlandiu arcticu Cray and other species in the Anvilian basin along the Chuckchee and Seward shores (Hopkins et al., 1974) and of planctonic foraminifera1 assemblages near the North Pole (Herman and Hopkins, 1980) and North Greenland shores (Funder et al., 1985). The climate during the l l t h , 9th and 7th thermochrons was almost equally warm; the mean winter temperature at that time in Western Europe was 1 to 2°C higher than at present (Van der Hammen et al., 1971) and in eastern Europe it was 5 to 10°C higher than at present. The average summer temperature was close to the modern level (Van der Hammen et al., 1971). The rainfall was higher than at present. The water temperature in the Tyrrhenian Sea was 0.5"C to 1.5"C higher than nowadays (Thunnell, 1979) both in summer and winter. SCT 6 is the great climatic boundary of the Late Cenozoic, the first continental ice-sheet glaciation of Europe, the maximum glaciation in North America (Easterbrook and Boellstroff, 1981), and also in Patagonia (Mercer, 1976). Its traces are

Fig. 11.18. Location of northern taiga fossil vegetation (Kap Kobenhaven, Meigen Islands and WorthPoint Island) in the sections of Kap Kobenhaven formation, approximate age 1.6- 1.8 Ma, and presentday spreading of these species. The solid line shows the modern northern boundary of forest spreading (after Funder et al., 1985).

KRESTOVKA

SIBERIA

SECTION--€.

PALEOMAGNETICS. FAUNAS. FORMATIONS,

A

VEGETATION FROM SHER iOTHERS. 1979

EUROPE

N O R T H AMERICA

LATER TORlNGlAN

-

RANCHOLABREAN I 1

EVENT

EARLIER T O R l N G l A N

-DISPERSAL

Sec. 7

-

DISPERSAL

RANCHOLABREAN I

EVENT

Mimomys ( M s a r i n i l

IRVINGTONIAN

II

LATE BlHARlAN

-

-

DISPERSAL

EVENT

IRVINGTONIAN I

EARLY h MIDDLE BlHARlAN

D. gulielmi~torquatus

D. ranldem, Microtus irst

-

B C I O ~ U SAllopharomys , DISPERSAL D. renidsns, Microtus

0% FOREST

P. compnalls r e w o r k e d P. hopkinri, Hypolagus

3.6% F O R E S T

Allophaiomyr cf pliocaenicus Cromeromyr intermedius P r e d k r o s t o n y i compitalis Clethrronomys c l . rulilus L w u s 5p.

F i r s t Dicrostonyx Mtcrotus, Allophaiomys EVENT

BLANCAN

VILLANY IAN

V

Ir s I Redicrostonyx iopklnsl. Synaptomys DISPERSAL

2 % FOREST

-

ltmomys ( M p a r r u s l

EVENT

as1 arge Hypolagus BLANCAN

-

IV

------ - LATE( RVILLAFRANCHIAN EBIELICE) KEY

sec.12

Hiatus in section NO

Pateomagnetics

Polarity unstated

2

BLANCAN

w I-

3

[ S E A S O N A L ICE

111 Synaptomys First

-

EVENT

DISPERSAL

u,

o>

'

N O

m

-

EARLY VILLAFRANCHIAN (ARONDELLI-TRIVERSA)

BLANCAN

II

-____-----

20% F O R E S T

------BLANCAN

I

CSARNOTAN

First Synaptomys (PuOCtOrnYSl REPENNING

Fig. 11.19. Correlation of Late Pliocene events in north-eastern Asia (section o n the Kresrovka river, a tributary of the Kolyma), North America and Europe (after Repenning, 1987).

w iD VI

396

found in eastern Europe and in the tills in cores in the town of Solikamsk on the Kama river and towns of Olonets and Vologda (Krasnov and Zarrina, 1986). In the middle parts of the Don basin it can be associated with boulders of erratic crystalline rocks in the lower Gorian beds, for which the Apsheronian ostracode assemblages and the mammal fauna of Odessa-Tamanian type are characteristic. A normal polarity zone is identified there which is referred to as the Olduvai (1.7 Ma) after Krasnenkov et al. (1984, 1987) and the Cobb Mountain (1.1 Ma) after the present authors. The erratic boulders show that in SCT 6 the ice sheet edge was at least in the Volga - Oka interfluve area. The moraines of the Menapian and the Narevian glaciations are found on the southern coast of the Baltic. Thus, the size of the ice sheet in SCT 6 was not smaller than the Pomeranian. Now it is clear that the replacement of the Villafranchian mammal fauna by that of the Pleistocene proper and the change in flora and vegetation in Georgia (Shatilova, 1974) between 1.1 and 1 .O Ma were caused by the glaciation of SCT 6. It is also significant that the Cassian regression which occurred simultaneously with that glaciation (Ambrosetti et al., 1972; Azzaroli, 1977, 1983) separates the two different states of the ocean level: before SCT 6 sea level was at height of about 25 m during all the thermochrons and after the 6th SCT it was only 1 - 5 m (Blackwelder, 1981). This means that beginning with SCT 6 the East Antarctic ice sheet was practically in a steady state and apart from the ice melting in the Northern Hemisphere only the West Antarctic ice sheet also retreated. Therefore, SCT 6 can be considered as the onset of the epoch of classical sheet glaciations in Europe, i.e. the onset of the Pleistocene in its traditional understanding (Zubakov, 1977, Zubakov et al., 1987).

11.5. Tentative reconstruction of climatic conditions for the Northern Hemisphere during the Middle Pliocene In this section, an attempt has been made to simulate climatic conditions (the temperature of the warmest and coldest months) for the Northern Hemisphere during the so-called “climatic optimum” of the Pliocene, which refers to a long period from SCT 25 to SCT 17 (zone N 19 Blow in marine section). According to the estimates of Budyko et al. (1980, 1982) during this period the atmosphere contained approximately twice as much carbon dioxide as at present, and it could be used as an analogue of the possible climate of the middle of the 21st century. On the basis of the paleogeographic, paleobotanical and paleontological data a preliminary sketch of the landscape of the Northern Hemisphere has been prepared to determine the expected air temperature of the warmest (July - August) and the coldest (January - February) months. The zonal method was used, which allowed us t o pass to climatic parameters of the past using the relationship between modern plant association and modern climate (Sinitsyn, 1965, 1967, 1980; Khotinsky and Savina, 1985; Liberman et al., 1985). For the ocean, the main information sources were data of factor and oxygenisotope analyses obtained in carrying out the deep-sea drilling project. Thus, the data of the oxygen-isotope analysis of the sites of the “Glomar Challenger” 5 5

397

(Douglas and Savin, 1971), 167 (Douglas and Savin, 1973), 187 and 189 (Keigwin, 1979), etc., located in the equatorial region, show that the surface sea-water temperature in the second half of zone N 19 by Blow (1969) and in zone NN 14 by Martini (1971), i.e. between 5.0 and 3.3 Ma ago, was 1.0- 1.5"C higher than at the present time. A pronounced thermal maximum in the equatorial regions is seen in the data of Van Goersel and Troelstra (1981) on the section of the Solo River in Java (Fig. 11.16), where throughout the entire zone N 19 the dextral form of Globorotalia menardii prevailed and the cold-water species of microfauna were practically absent. T o reconstruct the temperature field of surface layers of the ocean in temperate latitudes, the dependence of the present associations of planktonic foraminifera on surface water temperature has been used. According to the distribution of modern species of foraminifera, BC (1977) and Barash (1983) distinguish several latitudinal zones, the boundaries of which are closely connected with corresponding air and water temperatures. Still, Bandy et al. (1971) have shown that the positions of the planktonic foraminifera zones representing tropical temperate and subarctic waters considerably changed with latitude during the last 11 Ma. During the warm intervals the planktonic zones shifted noticeably northwards, by several degrees of latitude (up t o 10" at most), and during the coolings they moved towards the equator by approximately the same amount. New information, such as the detailed records of the factor analysis of Keller (1979) based on the DSDP sites 296, 310 and 173 from the Pacific, the data of Berggren (1978) based on the DSDP sites 111-A and 113 from the Atlantic, and the data of the oxygen-isotope analysis of the foraminifera shells from the subtropical and tropical regions of the Pacific and the Atlantic (Shackleton and Opdyke, 1977; Shackleton and Cita, 1979) allows us to reconstruct the temperature pattern of the ocean surface for the warmest (August) and the coldest (February) months of the year. The temperature difference between the Pliocene climatic optimum and the modern epoch is shown in Fig. 11.20 (a and b). The same maps present the temperature differences for the continents of the Northern Hemisphere obtained by data on paleofloras, including pollen records. Table 11.3 summarizes variations in the mean latitudinal temperature differences between the Pliocene optimum and modern time derived from maps and those obtained by Manabe and Bryan (1985) from the general circulation model with doubled CO, concentration. The paleoclimatic record as well as the model circulation show that the temperature change in middle and high latitudes was several times greater than that in lower latitudes. This fact is very important for the existence of the Arctic sea ice and West Antarctic ice sheet. According to Herman and Hopkins (1980) at the time of the Pliocene climatic optimum there was only seasonal ice in the Arctic basin, which completely disappeared in summer. According to Ciesielsky and Weaver (1974), the West Antarctic ice sheet degraded at that time as well. The greatest winter changes (January - February) are recorded north of 70"N in the Canadian archipelago region, Baffin Land and the north-east of the USSR, where the temperature increase was 20-22°C higher compared with the modern. In summer, the temperature change in high latitudes was also 6 - 8°C greater. No large areas

398

with negative values are observed in the regions, where precipitation considerably exceeded the modern (Kazakhstan, and the area now occupied by the Gobi and Sahara deserts). A greater area of small negative temperature departure was in North Africa, in the Sahara and the Sahel region, where at that time savanna and semi-savanna

Fig. 11.20. Deviation of winter (a) and summer (b) air temperature ("C) from modern one for the Pliocene optimum.

399

Table 11.3. Mean latitudinal temperature differences between the Pliocene climatic optimum and the modern epoch Latitude (day)

July August January February Annual mean Annual mean by Manabe and Bryan (1985) -

90-80

80-70

11.9

10.3

14.2 13.0

10.5

70-60

Global mean 0-90

60-50

50-40

40-30

30-20

6.5

5.4

4.1

2.3

1.1

13.2 11.8

11.3 8.9

6.4 5.9

4.8 4.4

3.8 3 .O

2.8 2.0

3.6

8.0

6.0

4.9

2.7

2.6

2.6

3.3

vegetation prevailed with richer flora and fauna than nowadays, which indicates sufficient moisture during the Pliocene maximum (Diester-Haas, 1979, Sarnthein et al., 1982). The change of mean annual global temperature obtained with regard to the areas of latitudinal zones by the palaeoclimatic data is 3.6"C compared with a value of 3.3"C according to Manabe and Bryan (1985). The relatively good agreement between these results shows that the period of the Pliocene maximum, i.e. 4.7 to 3.3 Ma ago, can be used as an analogue of the proposed climatic conditions in the mid21st century, when the C 0 2 concentration is expected to double.

RCsumC (1) Fifteen climatosedimentatary cycles have been identified in the undisturbed succession of coastal-marine sediments of the Pontean - Caspian and in terrestrial complexes of the southern Russian Plain in the period from the base of Pontian to the base of Chauda. The mean duration of these cycles ("zveno-cycle") has been estimated as 400 ka on the basis of the interpolation of paleomagnetic and oxygen isotipic records. This allowed us to reveal causal relationships between these cycles and the eccentricity rhythm. (2) Climatically opposite kryomer and thermomer parts of a zveno-cycle form a superclimathem, a unit convenient for the regional subdivision of Pliocene sediments. (3) Zveno-cycles and frequently superclimathems have more or less pronounced features of a biological (ecological) stratigraphic individuality. This makes it possible to consider a superclimathem as the major unit of the global correlation for the Pliocene. The Pontian -Caspian section can be suggested as their standard. The Mediterranean and North Sea sections have been studied with almost the same degree of detail as the Pontian - Caspian one.

400

(4) Global correlation of climatic events allows us to suggest three definitions for the concept of “thermal Pliocene optimum”: (a) In its broad conventional meaning it corresponds to the Brunsumian in Europe and foram zone N 19, that is covering an interval of 4.7 to 3.25 Ma (from SCT 23 to SCT 17); (b) In terms of paleoclimate it corresponds to hyperclimathem 8 marked in the sections by normal polarity events of Sidufial and Nunivak (4.67 - 4.09 Ma), that is it corresponds to the end of the Zanclian transgression and the Late Ruscinian in the Mediterranean Sea region, to the lower beds of the Sebastopol soil complex in the Ukraine, to the Cherlak formation in the Irtysh basin, the Khirgis-Nur (“B”) member in Mongolia, the lower beds of the Blancan in North America and nannoplankton zone N N 13. That was the time of sub-humid climate over the presentday deserts and semi-deserts, the time of rain forest fauna migration between Africa and Eurasia (Fig. 11.12). (c) Finally, in the narrow sense of paleooceanology, it correlates with superclimathem 21, that is the time of the highest Pliocene transgression (25 - 30 m) (Azores - Canary - Tabianian - Beringian - “Pekten”), K/Ar dated as 4.2 - 4.25 Ma when Pacific marine fauna penetrated to the Arctic basin and to some basins in East Antarctica (now ice-covered). (5) Paleoclimatic reconstructions carried out for the thermal Pliocene optimum in its broad sense (zones N 19 - the Brunsumian), 4.7 - 3.25 Ma ago, yield evidence indicating that the mean air temperatures for the Northern Hemisphere (0 - 80”) were equal to 3.6”C.

SUMMARY

The preparation of the English edition of the present publication has taken a lot o f time. When Part I had already been sent to Elsevier Publishers, the authors found two new publications relevant to the subject, namely Jenkins et al. (1985) and Sibrava et al. (1986) which give justification to the first Western versions of the general climatochronological scale for the Pleistocene. Both these scales by the principles of their development and by their content are similar to the scheme proposed by the authors (Zubakov and Borzenkova, 1983), presented in this book in a modified form. This important similarity, in our opinion, should increase the reader’s faith in the possibilities of climatostratigraphy. Obviously, certain differences in these schemes have been noticed. The main difference seems to concern the age estimates and the stratigraphic extent of the Likhvinian - Holsteinian - Hoxnian Interglacial sensu lato. According to one scheme (Jenkins et al., 1985; Sibrava et al., 1986; Zubakov and Borzenkova, 1983) the age corresponds to oxygen-isotopic stages 9 - 11; according to the other version, it corresponds to stages 9 - 11 - 13 (the present work). Apparently the problem needs to be further investigated. Equally uncertain remains the lower boundary of the Tyrrhenian - Karangatian - Eemian. The main conclusion of the three abovementioned publications seems to be that the Pleistocene “has matured” enough for detailed discussion of its general (international) climatostratigraphic (and chronological) scale. Certainly, this climathems subdivision of the Pliocene, not speaking of the Miocene, is still preliminary. First of all, it is due to doubtfulness of the age of the Messinian and Pontian. In many sections, the upper part of these regiostages represents erosional gap and the time hiatus possibly reaches 0.5 Ma. In this case, their upper boundary can have age of 4.85 Ma as can be observed on Crete and Apulia. Total duration of the Pontian increases, in this case, up to 2.3 Ma. These problems are debated in the author’s new book (V.A. Zubakov, Global Climatic Events of the Neogene) to be published by Hydrometeoizdat in 1990. We would like to introduce to the readers Part I1 of our book in this very context, though it presents only a small portion of the available factual data collected in the course of our studies. Our aim was to show that the pre-Pleistocene Cenozoic can also be a subject for a high-resolution climatostratigraphic correlation, and that the time is ripe to develop a paleoclimatic periodization for it. The differences between the Pleistocene, Pliocene and Miocene lie only in the duration of climathems, being recognized by paleontological and isotopic markers, and hence they can be traced globally. In the Pleistocene these are isotopic stages - orthoclimathems corresponding to parts of a 100 ka rhythm. In the Pliocene they are superclimathems, corresponding to parts of a 400 ka cycle, while in the Miocene they are hyperclimathems which are the parts of a 1200 ka cycle. We believe that the complete conventional stratigraphic scale, including the taxa of

402

the Standard Scale, zones of special biostratigraphic schemes, and also the regional stages and horizons of Regional Scales, all seem to have paleoclimatic causes. Being aware of this, i.e. after working out climato-chronological periodization independent of standard scales, we can increase the efficiency of our inter-regional stratigraphic correlations. And this is the first general inference from Part 11. The knowledge of the exact ages of past climatic events is the best tool in the understanding of the mechanisms and causes of climatic change. This is our second inference. Hence under the conditions of dramatically accelerating process of climatic change under the anthropogenic (technogenic) impact on the natural environment, when a reliable prediction of quantitative parameters of the climate for the near future becomes a necessary condition of the successful development of the world’s economy, the development of a reliable and high-resolution history of the past climates is of crucial importance for the present-day climatology. The idea of exospheric regimes and global climatic events, which seem to be determined by some external forcing and which can be empirically confirmed by the succession of climathems and their various taxa is but the first step to the bringing to light objective laws in the history of climate. Our schemes are far too schematic and formal. The authors nevertheless believe that their approach is rational. We shall appreciate opinions of our readers, even their sharp criticism.

ACKNOWLEDGEMENTS

In this book, there are a number of figures that were reproduced with the permission of the publishers, journals or authors. The authors of the book would like to thank: Publishers: American Geophysical Union, Washington (Figs. 2.1 1, 2.12, 8.3, 8.1 1, S.4 and S.6); National Academic Press, Washington D.C. (Figs. 1.3 and 10.9); Elsevier Science Publishers (Figs. 1.2, 2.1, 2.2, 3.2, 4.3, 8.1, and S.3); Geological Society of London (Figs. 10.1, 10.2 and 10.3). Editors of Initial Reports of Deep Sea Drilling Project (Figs. 2.3, 2.5, 2.7, 2.8 and 2.9). Editors of Journals: Nature (Figs. 8.9, S.5 and 11.10); Science (Figs. 8.2, 10.8 and 11.9); Marine Micropaleontology (Figs. 3.8, 10.5 and 11.16); Quaternary Research (Fig. 2.4); Boreas (Fig. 5.11), Meteor (Fig. 2.6); Episodes (Fig. 5.1); Geology (Fig. 11 .IS), Geologie en Mijnbouw (Fig. 5.3); Geologische Rundschau (Fig. 10.12) and Foraminifera1 Research (Fig. 11.17). Individuals: Prof. M.I. Budyko (Fig. 1.2); Dr. N. Chumakov (Figs. 1.5 and 1.6); Dr. E . Parisi and Prof. M.B. Cita (Fig. 3.1); Prof. V.L. Jakhimovich (Figs. 3.2 and 11.7); Prof. M.F. Veklich (Figs. 4.1 and 4.2); Dr. A.A. Lazarenko (Fig. 4.4), Prof. Y.E. Mojski (Fig. 5.2); Dr. Lindner (Fig. 5.4);Dr. M.P. Grichuk (Fig. 5.8); Prof. A. Dreimanis (Fig. 5.9); Prof. W.A. Berggren (Fig. 6.2); Dr. Yu. Vasilchuk (Fig. 8.10); Dr. H. Kessel and Prof. A. Raukas (Fig. 9.1); Dr. M.G. Grosswald (Fig. 9.4); Dr. T.N. Kaplina (Fig. 9.5); Drs. N.A., Jasamanov and A.N. Alexandrova (Figs. 10.6 and 10.7); Dr. H. Hooghiemstra (Fig. 11.1 l), and Dr. C.A. Repenning (Fig. 11.19).

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~

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SUBJECT INDEX

Anathermal - 256, 258, 276, 294 Bioclimatozone - 70, 75 Biological pump (pumping) - 270, 305, 307 - 3 I I Carbon dioxide (COz) concentration - 5 , 37, 269, 271, 301 in the air bubbles - 302 304 short-term variations - 302 - 3 I 1 Climate average (mean) climate - 6, 9, 15, 18, 38 definition - I S 16 geological climate - 15 local climate - IS Climathem (climatochron) 9, 10, 17, 18 19, 21, 23, 38, 297, 311, 324, 337 hyperclimathem - 23, 225, 337, 339, 342, 345, -

-

-

-

350, 384, 386, 401

-

-

352 - 354, 381 - 382, 384, 386, 399, 401 trendclimathem 23, 38, 325 -

Climatic cycle - see eccentricity cycle Climatic optimum see thermal optimum 5 Climatic palaeoanalog Climatic (exothermic) regime 10, 22, 28 glacial-psychrorpheric regime - 10, 28 29, 30, -

-

-

-

31, 37, 311, 325, 331

greenhouse-thermohaline regime

10, 28, 30,

-

306

-

205, 399-400 1200 ka cycle 337, 339, 350 3.70 Ma cycle 337, 339. 350 1.1 Ma cycle 337, 339, 350 Ebent (signal) stratigraphy 192, 218, 309. 311, 324 Forcing factors 253 Geomagnetic excurSions 58- 61, 62 - 63, 84, 110, 113-114. 116-117 Glacial-interglacial cycle - 42 Gunz 42, 199-200, 209-211 Mindel - 42, 199, 200, 212-215 Ris\ - 42, 199-200, 215-216, 219-223, 225 - 226 Wurm - 42, 199-200, 231, 232-235. 240-241, 244, 249 Global climatic event - 17 - 18, 21, 3 8 - 39. 93. 95 time-classification (periodization) - 2 I - 23, 198 199, 205 Katathermal - 250, 267, 272 -273. 276. 282, 294 Kryomer (kryochron) - 19, 22, 40, 205. 207, 251, 260, 297, 307, 3 1 1 Little Ice Age 304 Magnethem (magnetozone) 21 Marine terraces 223, 227, 244 Xlilankovitch theory - 298 Megathermal - 247, 256, 260, 267, 272 - 273, 276, 283, 287 -288, 294 -

-

-

-

-

-

-

-

-

Climatostratigraphic correlation - 18, 42, 54, 189 long distant correlation - 95, 189- 191, 297 synchroneity/metochroneity problem 194- 197, 207 Climatostratigraphy - 16, 17, 18, 40, 42 Clirnatostratigraphy scheme - see regional scheme Climatography - 16 Dating method\ - 19 geochronometric methods - 18 19, 88, 182 geomagnetic (paleomagnetic) - 19, 21, 42, 54, -

-

-

-

32-33, 35, 3 6 - 3 7 , 311, 325

88

-

-

nannoclimatheni - 254-255, 256, 294, 297, 299, 307 orthoclimathem 23, 54, 66, 205, 207, 297, 299, 307, 311, 401 superclimathem 205, 206, 298, 307, 31 I , 337,

El-Nitio

Eccentricity cycle, rhythm - see solar insolation change 100 ka cycle 30, 43 200 ka cycle 205 370- 413 ka cycle 200, 201, 202 - 203, 204,

%loislure conditions middle latitudes - 275 - 276, 286. 2Y I subtropical regions - 230, 276, 279 280. 284, -

291, 293

Oxygen-isotope 30, 43, 45-46, 4 8 - 5 3 , 56, 72. 3 0 0 , records -

326, 335, 341 stages - 43, 45, 48 - 52. 54, 207 Paleoclimate - 5 -6, 15 definition - 15 - 16

456 Paleoclimatic reconstructions - 4, 5, 45, 228 - 23 1, 246 - 247, 249, 266, 267, 275, 289, 297 Paleoclimatology - 5-7, 16 paradigms - 5-7, 10 Paleolithic tools - 110-111, 116-117, 119 Pleistocene climatostratigraphic classification 44 rhythm-chronological approach - 197 - 198, 200 Pliocene-Pleistocene boundary - 192 - 194, 380 Precipitation trend - 24, 120, 129, 159, 274, 276, 280, 286, 288-289, 353, 357 Psychrosphere - 27, 30 Rainfall annual - 231, 248, 280, 282, 289 monsoon - 283, 291, 294 Regional climatostratigraphic sequences (schemes, tables) Pleistocene: Alaska - 157 Arctic and Subarctic - 174 Black Sea Region - 84, 85 Caspian - 80, 82 Central Asia and China - I16 Mediterranean - 70, 72, 75 Middle Europe - 103 North America - 166, 174 North-Eastern Asia - 157 North-Western Europe - 133 Poland - 133 Russian Plain - 127 West Siberian - 146 Pliocene: Black Sea - 352

Caspian - 358 Mediterranean - 375 West Siberian - Kazakhstan - 383 Ocean level changes - 385 Miocene: Tethys-Paratethys - 338, 339 Sapropel layers - 72, 95 - 28, 236 Sea ice - 310, 380 Soil - 103, 106, 111-112, 116-117, 120 Shumer drift - 299 Solar insolation changes due to orbital factor (Milankovitch theory) - 51, 269, 271 -272, 274, 292 - 294, 298, 300, 305 Surges, surging - 260, 299, 308 Speleothems - 170 Termination - 301, 302, 307 Thermal optimum Eocene optimum - 33 - 37, 328 Holocene optimum - 266 - 267 Mikulino optimum - 227 - 231 Miocene optimum - 335, 339, 341 - 343, 346 - 349 Pliocene optimum - 388, 390, 396-400 Thermomer (thermochron) - 19, 22, 40, 205, 207, 251, 254, 256, 258, 267-268, 297, 307, 311, 382 Temperature trend - 24, 120, 129, 136, 139, 149 159, 224, 238, 253, 256, 259, 271, 212, 291, 294, 301, 357 Upwelling - 306, 308 - 309, 341, 344 Volcanic aclivity - 37, 268, 272, 274, 294, 300-301, 307-308, 311 Zveno-cycle - 399 - 400