Geology and Geoarchaeology of the Black Sea Region: Beyond the Flood Hypothesis (GSA Special Paper 473)

Geology and Geoarchaeology of the Black Sea Region: Beyond the Flood Hypothesis edited by Ilya V. Buynevich Department ...

Author: Ilya V. Buynevich | Valentina Yanko-Hombach | Allan S. Gilbert | Ronald E. Martin

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Geology and Geoarchaeology of the Black Sea Region: Beyond the Flood Hypothesis edited by Ilya V. Buynevich Department of Earth and Environmental Science Temple University 1901 N. 13th Street Philadelphia, Pennsylvania 19122, USA Valentina Yanko-Hombach Avalon Institute of Applied Science 976 Elgin Avenue Winnipeg, Manitoba R3E 1B4, Canada and Department of Physical and Marine Geology Odessa National I.I. Mechnikov University 2 Shampansky per. Odessa, 65058, Ukraine Allan S. Gilbert Department of Sociology and Anthropology Dealy Hall 401 Fordham University Bronx, New York 10458, USA Ronald E. Martin Department of Geological Sciences University of Delaware 103 Penny Hall Newark, Delaware 19716-2544, USA

Special Paper 473 3300 Penrose Place, P.O. Box 9140

Boulder, Colorado 80301-9140 USA

2011

Copyright © 2011, The Geological Society of America (GSA), Inc. All rights reserved. GSA grants permission to individual scientists to make unlimited photocopies of one or more items from this volume for noncommercial purposes advancing science or education, including classroom use. For permission to make photocopies of any item in this volume for other noncommercial, nonprofit purposes, contact The Geological Society of America. Written permission is required from GSA for all other forms of capture or reproduction of any item in the volume including, but not limited to, all types of electronic or digital scanning or other digital or manual transformation of articles or any portion thereof, such as abstracts, into computer-readable and/or transmittable form for personal or corporate use, either noncommercial or commercial, for-profit or otherwise. Send permission requests to GSA Copyright Permissions, 3300 Penrose Place, P.O. Box 9140, Boulder, Colorado 80301-9140, USA. GSA provides this and other forums for the presentation of diverse opinions and positions by scientists worldwide, regardless of their race, citizenship, gender, religion, or political viewpoint. Opinions presented in this publication do not reflect official positions of the Society. Copyright is not claimed on any material prepared wholly by government employees within the scope of their employment. Published by The Geological Society of America, Inc. 3300 Penrose Place, P.O. Box 9140, Boulder, Colorado 80301-9140, USA www.geosociety.org Printed in U.S.A. GSA Books Science Editors: Marion E. Bickford and Donald I. Siegel Library of Congress Cataloging-in-Publication Data Geology and geoarchaeology of the Black Sea Region : beyond the flood hypothesis / edited by Ilya V. Buynevich ... [et al.]. p. cm. -- (Special paper ; 473) Includes bibliographical references. ISBN 978-0-8137-2473-7 (pbk.) 1. Geology--Black Sea Region. 2. Archaeological geology--Black Sea Region. 3. Paleoclimatology-Holocene. I. Buynevich, Ilya V. (Ilya Val) QE350.22.B55G465 2011 554.9--dc22 2010046616 Cover: Satellite image of the Black Sea. NASA image courtesy of the MODIS Rapid Response Team (http://earthobservatory.nasa.gov/IOTD/view.php?id=8817). Inset, left: Eroding cliffs of Berezan Island, Ukraine, an important archaeological site along the northern Black Sea coast. Photo by I. Buynevich. Inset, center: ROV Hercules over a shipwreck with amphorae on the bottom of the Black Sea. Photo ©IFE/COE. Inset, right: Remnants of Tauric Chersonesos, an important Greek colony and port on the Crimean Peninsula. Photo by I. Buynevich.

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Contents

Preface . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . v V. Yanko-Hombach, A.S. Gilbert, I.V. Buynevich, and R.E. Martin 1. Surface runoff to the Black Sea from the East European Plain during Last Glaciation Maximum–Late Glacial time . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1 A.Yu. Sidorchuk, A.V. Panin, and O.K. Borisova 2. Modeling extreme Black Sea and Caspian Sea levels of the past 21,000 years with general circulation models . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 27 A. Kislov and P. Toropov 3. Assessment of the Black Sea water-level fluctuations since the Last Glacial Maximum . . . . . . . 33 G. Lericolais, F. Guichard, C. Morigi, I. Popescu, C. Bulois, H. Gillet, and W.B.F. Ryan 4. Rapid Holocene sea-level and climate change in the Black Sea: An evaluation of the Balabanov sea-level curve . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 51 R.E. Martin and V. Yanko-Hombach 5. Global climate change and sea-level fluctuations in the Black and Caspian Seas over the past 200 years . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 59 E. Konikov and O. Likhodedova 6. Paleogeography of the Pontic Lowland and northwestern Black Sea shelf for the past 25 k.y. . . 71 E. Larchenkov and S. Kadurin 7. Nonpollen palynomorphs: Indicators of salinity and environmental change in the Caspian–Black Sea–Mediterranean corridor . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 89 P.J. Mudie, S.A.G. Leroy, F. Marret, N.P. Gerasimenko, S.E.A. Kholeif, T. Sapelko, and M. Filipova-Marinova 8. Climatic and environmental oscillations in southeastern Ukraine from 30 to 10 ka, inferred from pollen and lithopedology . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 117 N.P. Gerasimenko 9. Late Pleistocene and Holocene paleoenvironments of Crimea: Pollen, soils, geomorphology, and geoarchaeology . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 133 C.E. Cordova, N.P. Gerasimenko, P.H. Lehman, and A.A. Kliukin

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10. Bedforms, coastal-trapped waves, and scour process observations from the continental shelf of the northern Black Sea . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 165 A. Trembanis, S. Nebel, A. Skarke, D.F. Coleman, R.D. Ballard, A. Yankovsky, I.V. Buynevich, and S. Voronov 11. Archaeological oceanography and environmental characterization of shipwrecks in the Black Sea . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 179 M.L. Brennan, R.D. Ballard, K.L. Croff Bell, and D. Piechota 12. Pontic-Baltic pathways for invasive aquatic species: Geoarchaeological implications. . . . . . . . 189 I.V. Buynevich, A. Damušytė, A. Bitinas, S. Olenin, J. Mažeika, and R. Petrošius

Preface These opening words convey only a few essential matters. The present volume is one of a growing number of works focusing on the Black Sea, and as such, its relationship to previous research and its links to that of the near future beg some clarification, and thereby perspective. In addition, no interdisciplinary publication is ever achieved without help from a wide range of contributors whose part in the process deserves a public statement of deep appreciation. The Black Sea is the largest anoxic basin in the world, encompassing a total area of 423,000 km2. The basin is surrounded by Alpide fold belts and was formed in the Mesozoic as a backarc structure above the northward-subducting Tethyan oceanic lithosphere. The Black Sea consists of two large subbasins on the west and east that are separated by the NW-SE–trending Mid–Black Sea ridge. The western subbasin is floored by oceanic crust over which lie thick sediment units probably of Cretaceous and younger age; the eastern subbasin has a thinned continental or oceanic crust with a sediment cover less than 10 km thick. As a marginal basin, the Black Sea acts as a paleoenvironmental amplifier, recording climatic events in great detail. In response to sea-level changes driven by climatic cycles and/or regional tectonics, its connections with adjacent basins (the Marmara, Mediterranean, and Caspian Seas) have periodically been altered, leading to coastline migration and drastic modifications in environmental conditions (i.e., salinity, oxygen regime, basin morphology, hydrology), with dramatic consequences for the sedimentary, geochemical, and ecological systems, as well as human adaptive strategies. RENEWED SCIENTIFIC INTEREST IN THE BLACK SEA REGION Lately, this basin has witnessed a tremendous surge in interest due to (1) the Great Flood hypotheses that tied the Biblical Flood to the Black Sea (Ryan et al., 1997, 2003; Chepalyga 2003, 2007), (2) the presence of huge methane reserves contained within gas hydrates beneath the seafloor that may be exploitable as new nontraditional energy sources (Shnyukov and Ziborov, 2004), (3) the growing tangle of underwater infrastructure (e.g., gas pipelines and communication cables) laid across the Black Sea floor that is increasingly subject to geohazards from landslides, tectonics, and other dynamic forces, and (4) the presence of vast amounts of raw materials (e.g., sapropels) that have economic applications in agriculture (Shnyukov et al., 1999). This new outlook on the Black Sea has fostered a series of meetings, symposia, and workshops targeting issues in the geology, climatology, geochemistry, and archaeology of the Pontic basin. Three of them, held in 2003—(1) NATO Advanced Research Workshop “Climate Change and Coastline Migration,” 1–5 October 2003, in Bucharest, Romania; (2) international conference “The Black Sea Flood: Archaeological and Geological Evidence” sponsored by the Columbia University Seminar on the Ancient Near East, 18–20 October 2003, in New York, USA; and (3) Geological Society of America (GSA) Topical Session “‘Noah’s Flood’ and the Late Quaternary Geological and Archaeological History of the Black Sea and Adjacent Basins” presented at the Geological Society of America Annual Meeting on 4 November 2003, in Seattle, USA—led to the publication of a 1000-page volume entitled The Black Sea Flood Question: Changes in Coastline, Climate, and Human Settlement, which appeared under the Springer imprint (Yanko-Hombach et al., 2007a). This volume included 35 papers dealing with the geological, hydrological, climatological, archaeological, and linguistic aspects of the Black Sea flood hypotheses. Although no final answer to the Black Sea flood question appeared there, the book made great strides in enabling expanded dialogue between western and eastern scientists, encouraging new collaborations, and familiarizing western researchers with the extensive amount of information obtained by eastern scientists, data that had previously been inaccessible owing to the local languages in which they had originally been published. Subsequently, east-west collaboration continued to grow in the research programs of individual scientists as well as in international multidisciplinary projects, such as International Geological Correlation Programme (IGCP) 521 “The Black Sea–Mediterranean Corridor during the last 30 k.y.: Sea-level change and human adaptation” and International Union for Quaternary Research (INQUA) 501 “The Caspian–Black v

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Lithuania Denmark Russia Canada USA

UK Ireland France Romania Bulgaria

Ukraine

Egypt

Black Sea

Figure 1. International Geological Correlation Programme (IGCP) 521 and International Union for Quaternary Research (INQUA) project 501 participating countries (highlighted on the map): Algeria, Australia, Austria, Azerbaijan, Belgium, Bulgaria, Canada, Croatia, Egypt, Germany, Finland, France, FYR of Macedonia, Georgia, Greece, Ireland, Israel, Italy, Kazakhstan, Latvia, Lithuania, Moldova, Romania, Russian Federation, Spain, Switzerland, The Netherlands, Turkey, Ukraine, UK, United States of America (countries with contributors to this volume are listed on the map). Note the wide geographic distribution of scientists carrying out research in the Black Sea region.

Sea–Mediterranean Corridor during the last 30 k.y.: Sea-level change and human adaptive strategies” (www .avalon-institute.org/IGCP). Today, these projects involve the work of ~400 scientists, not only from the Black Sea region, but from around the world (Fig. 1). After the first three conferences in 2003, five plenary meetings were conducted under the framework of the IGCP 521–INQUA 501 projects from 2005 to 2009, and numerous topical sessions were presented at leading geological forums, such as the Annual Assembly of the European Geological Union in Vienna, Austria (2005, 2006); the Annual Meeting of the Geological Society of America in Denver, USA (2007); the XIIth INQUA Congress in Cairns, Australia (2007); and the 33rd International Geological Congress in Oslo, Norway (2008). In addition, many other smaller meetings examined in further detail the flood hypotheses but also addressed issues of regional climate, tectonics, coastline migration, human adaptive strategies, economic resources, and the future environmental stability of the region. More than 1000 authors have made contributions to IGCP 521–INQUA 501 meetings by presenting ~500 papers (Yanko-Hombach et al., 2005, 2006, 2007b; Gilbert and Yanko-Hombach, 2008, 2009). Many of these papers have been published or will be published in five IGCP 521–INQUA 501 thematic volumes of Quaternary International. THE DENVER CONFERENCE AND COMPILATION OF THE VOLUME The papers contained within this special GSA volume are the outgrowth of a successful technical session at the 2007 Geological Society of America Annual Meeting in Denver, Colorado. A large number of participants from Eastern Europe, funded by the GSA International Division, had the opportunity to present their recent findings, and their contributions are an integral part of the volume. The twelve papers were written by contributors from twelve countries (Fig. 1), and they address a range of topics, including climatic and hydrologic modeling, paleogeographic reconstruction of late Quaternary landscapes, palynology and paleoclimate reconstruction, and geoarchaeological studies, both onshore and offshore. We hope that the volume will serve as a timely reference for continuing research in a region harboring a number of newly independent states that are now faced with population pressure and a variety of environmental issues.

Preface

Each paper in the present book underwent a lengthy review process (three reviewers as a rule per paper) and both language and graphics editing. Acknowledgment must first be given for the financial assistance that made the conferences and book possible. We thank the International Union of Geological Sciences (IUGS), IGCP, INQUA, United Nations Educational, Scientific, and Cultural Organization (UNESCO), and GSA, which provided grant sponsorship to support many presenters at IGCP 521– INQUA 501 meetings. All transliterations of cited sources in Cyrillic follow Library of Congress style for both consistency and compatibility with the Online Computer Library Center’s World Catalogue, to maximize ease of location for the references in question. Grateful acknowledgment is offered for the thoughtful efforts of many external reviewers: Patrick Conaghan, Australia; Veselin Peychev, Bulgaria; John McAndrews, Petra Mudie (internal editorial help), Canada; K. Petersen, Denmark; Goran Georgievski, Jürgen Herget, Jens Matthiessen, Germany; Eliso Kvavadze, Georgia; Michel Fontugne, France; Daniella Basso, Italy; Tomasz Kalicki, Poland; Oya Algan, Mustafa Ergin, Namık Çağatay, Erdinç Yiğitbaş, Turkey; Doug Levin, Antonio Rodriguez, Shelley Wachsmann, USA, and a number of anonymous reviewers. Lastly, we thank the GSA book editor’s staff for, above all, their patience in awaiting the delivery of the finished manuscript. Valentina Yanko-Hombach Co-Leader of IGCP 521 and Leader of INQUA 501 Allan S. Gilbert Ilya V. Buynevich Ronald E. Martin

REFERENCES CITED Chepalyga, A.L., 2003, Late Glacial Great Flood in the Black Sea and Caspian Sea: Geological Society of America Abstracts with Programs, v. 35, no. 6, p. 460. Chepalyga, A.L., 2007, The Late Glacial Great Flood in the Ponto-Caspian basin, in Yanko-Hombach, V., Gilbert, A.S., Panin, N., and Dolukhanov, P., eds., The Black Sea Flood Question: Changes in Coastline, Climate and Human Settlement: Dordrecht, Springer, p. 119–148. Gilbert, A., and Yanko-Hombach, V., eds., 2008, Extended Abstracts of the Fourth Plenary Meeting and Field Trip of IGCP 521–INQUA 501 Project “Black Sea–Mediterranean Corridor during the Last 30 k.y.: Sea Level Change and Human Adaptation,” 4–16 October 200: Bucharest, Romania, National Institute of Marine Geology and Geoecology (GeoEcoMar), and Varna, Bulgaria, Department of Natural History of the Regional Historical Museum, 215 p., ISBN 978-973-0-06271-7. Gilbert, A., and Yanko-Hombach, V., eds., 2009, Extended Abstracts of the Fourth Plenary Meeting and Field Trip of IGCP 521–INQUA 501 Project “Black Sea–Mediterranean Corridor during the Last 30 k.y.: Sea Level Change and Human Adaptation,” 22–31 August 2009: Izmir, Turkey, Kadir Has University, Dokuz Eylül University, and Çanakkale, Turkey, Çanakkale Onsekiz Mart University, 213 p., ISBN 978-975-441-265-9. Ryan, W.B.F., Pitman, W.C., III, Major, C.O., Shimkus, K., Maskalenko, V., Jones, G.A., Dimitrov, P., Görür, N., Sakinç, M., and Yüce, H., 1997, An abrupt drowning of the Black Sea shelf: Marine Geology, v. 138, p. 119–126, doi: 10.1016/S0025-3227(97)00007-8. Ryan, W.B.F., Major, C.O., Lericolais, G., and Goldstein, S.L., 2003, Catastrophic flooding of the Black Sea: Annual Review of Earth and Planetary Sciences, v. 31, p. 525– 554, doi: 10.1146/annurev.earth.31.100901.141249.

Shnyukov, E., and Ziborov, A., 2004, Mineral’nie bogatstva Chernogo moria [Mineral Riches of the Black Sea]: Kiev, Department of Marine Geology and Mineral Resources of the Ukrainian Academy of Sciences. Shnyukov, E.F., Kleschenko, S.A., and Kukovskaya, T.S., 1999, Sapropelevie ili Chernogo moria—Novii vid mineral’nogo siriia [Sapropels of the Black Sea—New kind of raw materials]. Geologiia i poleznie iskopaemie Chernogo moria [Geology and Mineral Resources of the Black Sea]: Kiev, p. 399–411. Yanko-Hombach, V., Buynevich, I., Chivas, A., Gilbert, A., Martin, R., and Mudie, P., eds., 2005, Extended Abstracts of the First Plenary Meeting and Field Trip of IGCP 521 Project “Black Sea–Mediterranean Corridor during the Last 30 k.y.: Sea level Change and Human Adaptation,” 8–15 October 2005: Istanbul, Turkey, Kadir Has University, 226 p. Yanko-Hombach, V., Buynevich, I., Chivas, A., Gilbert, A., Martin, R., and Mudie, P., eds., 2006, Extended Abstracts of the Second Plenary Meeting and Field Trip of IGCP 521 Project “Black Sea–Mediterranean Corridor during the Last 30 k.y.: Sea Level Change and Human Adaptation,” 20–28 August 2006: Odessa, Ukraine, Odessa National University, 188 p., ISBN 966-318-554-6. Yanko-Hombach, V., Gilbert, A.S., Panin, N., and Dolukhanov, P.M., eds., 2007a, The Black Sea Flood Question: Changes in Coastline, Climate and Human Settlement: Dordrecht, The Netherlands, Springer, 971 p. Yanko-Hombach, V., Buynevich, I., Dolukhanov, P., Gilbert, A., Martin, R., McGann, M., and Mudie, P., eds., 2007b, Extended Abstracts of the Joint Plenary Meeting and Field Trip of IGCP 521 “Black Sea–Mediterranean Corridor during the Last 30 k.y.: Sea Level Change and Human Adaptation,” and IGCP 481 “Dating Caspian Sea Level Change,” 8–17 September 2007: Gelendzhik (Russia)-Kerch (Ukraine), Southern Branch of the Institute of Oceanology, Russian Academy of Sciences and Demetra Beneficent Foundation, 178 p., ISBN 978-5-85941-151-0, IGCP 521-481.

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The Geological Society of America Special Paper 473 2011

Surface runoff to the Black Sea from the East European Plain during Last Glacial Maximum–Late Glacial time Aleksey Yu. Sidorchuk Andrey V. Panin Geographical Faculty, Moscow State University, Vorob’evy Gory, Moscow 119991, Russia Olga K. Borisova Institute of Geography, Russian Academy of Sciences, Staromonetny per., Moscow 119017, Russia

ABSTRACT Hydromorphological and hydroclimatic methods were used to reconstruct the former surface runoff from the East European part of the Black Sea drainage basin. Data on the shape and dynamics of the last Fennoscandian ice sheet were used to calculate meltwater supply to the headwaters of the Dnieper River. The channel width and meander wavelength of well-preserved fragments of large paleochannels were measured at 51 locations in the Dnieper and Don River basins (East European Plain), which allowed reconstruction of the former surface runoff of the ancient rivers, as well as the total volume of flow into the Black Sea, using transform functions. Studies of the composition of fossil floras derived from radiocarbon-dated sediments of various origins and ages make it possible to locate their modern region analogues. These analogues provide climatic and hydrological indexes for the Late Pleniglacial and Late Glacial landscapes. Morphological, geological, geochronological, and palynological studies show that the landscape, climatic, and hydrologic history of the region included: (1) a cold and dry interval close to the Last Glacial Maximum characterized by high meteoritic surface runoff supplemented by meltwater flow from ice-dam lakes; (2) a warmer humid interval at the end of the Late Pleniglacial with very high surface runoff and formation of extremely large meandering channels, combined with a short event of substantial inflow from the Caspian Sea; and (3) a period from the Oldest Dryas to the Preboreal of nonsteady surface runoff decrease, and transformation of large meandering channels into smaller ones against the background of climate warming.

Sidorchuk, A.Yu., Panin, A.V., and Borisova, O.K., 2011, Surface runoff to the Black Sea from the East European Plain during Last Glacial Maximum–Late Glacial time, in Buynevich, I.V., Yanko-Hombach, V., Gilbert, A.S., and Martin, R.E., eds., Geology and Geoarchaeology of the Black Sea Region: Beyond the Flood Hypothesis: Geological Society of America Special Paper 473, p. 1–25, doi: 10.1130/2011.2473(01). For permission to copy, contact [email protected]. © 2011 The Geological Society of America. All rights reserved.

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INTRODUCTION General information on the variability of climate and water budget in Europe for the period from the Last Glacial Maximum (LGM), through Late Glacial time (LGT), and into the beginning of the Holocene (ca. 18–10 radiocarbon ka) is controversial. According to vegetation reconstructions based on palynological data (Grichuk, 1982), the climate is believed to have been both cold and dry (the so-called cryoxerotic stage of the glaciation). The climate of the southern part of the East European Plain drives changes in the water balance of southern sea and lake basins, and, thus, most workers correlate the last major drop in the level of the Caspian Sea and Black Sea to the LGM (Varuschenko et al., 1987; Winguth et al., 2000; Dolukhanov et al., 2008; and others). To explain such a dramatic drop in sea level, a substantial decrease in river runoff into the seas has been suggested. Varuschenko et al. (1987) estimated annual river runoff into the Caspian Sea during the LGM at only 20%–28% of its present value. Estimates of river runoff into the Caspian and Black Seas during the LGM based on atmospheric general circulation models comprise 55% and 59% of the modern values, respectively (Kislov and Toropov, 2006). On the other hand, extensive dating of the Caspian deposits undertaken in the last decade have revealed a Late Glacial age for the highest stage of the Khvalynian transgression (Svitoch and Yanina, 1997; Leonov et al., 2002; Chepalyga et al., 2008; and others), which suggests large surface runoff, at least from the Volga River basin. Kalinin et al. (1966) estimated river runoff into the Caspian Sea during the maximum stage of the Khvalynian transgression at 517 km3 per year, a figure that is 1.5 times higher than the present value. Similarly controversial reconstructions exist for the Black Sea drainage basin: some researchers propose a relatively high runoff and a continuous outflow from the Black Sea (Lane-Serff et al., 1997), while others believe that the runoff into the Black Sea was low due to the dry climate, and as a result, sea level dropped to –110 m (Ostrovsky, 1982; Aksu et al., 2002) or even to –150 m (Winguth et al., 2000) because of negative water budget during the LGM. It has also been suggested that a massive inflow of meltwater from the Fennoscandian ice sheet into the Black Sea took place after the LGM (Kvasov, 1979; Kroonenberg et al., 1997), both directly through the Dnieper River valley and as an outflow from the Caspian Sea through the Manych Straight. A series of meltwater pulses is suggested by isotopic depletion of the Black Sea waters between 18 and 15.5 ka (Bahr et al., 2006). Hydromorphological and hydroclimatic methods of paleogeographic reconstructions allow quantitative estimation of the former surface runoff originating from melting glaciers and from precipitation over a river basin. The ice volume of the Quaternary ice sheets can be reconstructed from their area using transform functions derived from recent glaciological information and theoretical considerations about ice rheology (Markov and Suetova, 1964; Khodakov, 1982; Peltier, 1994). Therefore, with information on the age of the boundaries of the last Fennoscandian ice

sheet, changes in its volume and meltwater supply into adjacent rivers can be estimated (Kalinin et al., 1966). Valuable information about past hydrological river regimes can be derived from the morphology of the former river channels, especially if the former topography differs significantly from that of the present. Morphological data on the large Late Glacial paleorivers, which give distinct evidence of high surface runoff, were first investigated by Dury (1964, 1965) in Western Europe and North America, and by Volkov (1960, 1963) in northern Kazakhstan and western Siberia. Similar results were obtained for several rivers in the Black Sea basin: for the Seim (Borisova et al., 2006) and Khoper (Sidorchuk and Borisova, 2000) Rivers in the Dnieper and Don basins; for the basin of the Danube River in Hungary (Borsy and Felegyhazi, 1983; Kasse et al., 2000), and in Romania (Howard et al., 2004). Morphological, textural, palynological, and geochronological studies have shown that the LGT in Europe was a period of very large, widely spread river channels, which presumably were formed by high and powerful surface runoff. Paleobotany plays an important role in providing data for paleoclimatic reconstructions. Late Glacial climatic events and the chronology of vegetation development in Europe have been derived mainly from palynological data later dated through radiocarbon and correlated with the isotopic “events” in the Greenland ice-core record (e.g., Walker et al., 1999). To reconstruct the hydroclimatic conditions that existed at various stages of the LGM and LGT, palynological studies of dated alluvium, lake, and peat sediments can be applied. The use of paleobotanic data for paleoclimatic and paleolandscape reconstruction implies that flora and vegetation are strongly influenced by changes in the natural environment and by the climate in particular (e.g., Iversen, 1944; Grichuk, 1969). This paper is aimed at (1) analysis of the paleoenvironmental conditions and causes for the paleohydrological changes in the East European part of the Black Sea basin during the LGM and LGT, (2) paleohydrological reconstruction of the surface runoff there since the Late Pleniglacial, and (3) discussion of the landscape and surface runoff changes in the remaining part of the Black Sea basin during the last ~20,000 yr. METHODS Paleofloristic Method of Paleolandscape and Hydroclimatic Reconstruction Climate reconstructions usually rely on either the comparison of fossil and modern pollen assemblages and their associated modern climate, or selected indicator plant species with specific climatic requirements. Paleobotanical data used for such reconstructions are of two main kinds: plant macrofossils (seeds, fruits, leaves, etc.), and pollen and spores. Macrofossils have the advantage of usually being identifiable to the species level. On the other hand, the occurrence of macrofossils is relatively restricted, and they usually belong to aquatic and subaquatic

Surface runoff to the Black Sea from the East European Plain plants. Pollen diagrams provide more detailed information on the history of specific plant taxa as well as vegetation on the whole. Because plants sensitive to climatic conditions are mainly medium to low pollen producers, sufficiently detailed pollen data are essential. In the process of pollen identification, the highest possible taxonomic resolution should be achieved to obtain the most complete results. This is possible if well-preserved pollen of arboreal plants, as well as pollen and spores of certain groups of herbaceous plants (e.g., Thalictrum, Lycopodium, Equisetum), can be identified to genus or even species levels. Iversen (1944) was the first to use the pollen of certain plant species to estimate paleotemperatures. This author established the relationship between present occurrences of Ilex and Hedera and summer as well as winter temperatures. Such relationships can be established by comparing the present boundaries of the plant’s geographical range (its “area”) and climatic data, i.e., the coincidence of “area” limits and certain isotherms. For example, Iversen showed that present-day ranges of Ilex and Hedera are limited by the 0 °C and –2 °C January isotherms, respectively. A well-known example of this kind is the coincidence of the tree line with the 10 °C July isotherm in many lowland and mountain areas. This approach has subsequently been extended by increasing the number of indicator species and climatic indexes considered (e.g., Hintikka, 1963; Zagwijn, 1996). Grichuk (1969) further developed Iversen’s approach into the mutual climatic range method, in which as many species as possible are used to obtain paleoclimatic estimates for a specific site. Their geographical distributions are first converted to climatic ranges, expressed as climatic range diagrams (climatograms), and then their mutual climatic field is determined. Climatograms have the advantage that even plant species without present-day geographical overlap can be used for paleoclimatic reconstructions. Moreover, since this approach is based on the requirements of individual species, finding modern analogues for the fossil pollen spectra or plant assemblages is not necessary. Grichuk (1969) suggested yet another method of climatic reconstruction based on paleofloristic data, elaborating on the idea of Szafer (1946–1947). This method consists of identifying a modern region where all the species of a fossil flora grow presently. The mutual area of all individual species of a certain fossil flora is found by overlapping their present-day “areas.” The “areas” of as many plants as possible should be used. Geographical analysis of the modern spatial distribution of all the plants of a certain fossil flora (compilation of a so-called arealogram) allows one to find the closest modern floristic analogue to the past vegetation at the site. By identifying this modern regional analogue, it is possible to determine the closest modern landscape and hydroclimatic environment to that of the fossil flora under study. Accuracy in these reconstructions depends on (1) the accuracy of the paleofloristic definitions based on detailed pollen analysis, (2) the richness of the resulting fossil floras, (3) the accuracy of the data on present-day geographical ranges of plants that represent components of the fossil floras, (4) the sizes of the regional analogues, and (5) the variability of the hydroclimatic

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characteristics within these analogues. Furthermore, the accuracy of palynological analysis and the richness of a paleoflora depend not only on the palynologist’s personal experience, but also on the type of sediment, vegetation type, and the degree of pollen preservation. Late Glacial floras are usually relatively poor, so that not all analyzed samples provide sufficient data for paleofloristic reconstruction. Usually, conditions that are suitable for all the species of a given fossil flora can be found within a comparatively small area. The present-day features of plant communities and the main hydroclimatic indices of such a regional analogue would be close, in most cases, if not identical, to those that existed at the sampled site in the past. For example, fossil flora from the cultural layer at the Yudinovo Early Man site in the Desna River basin consists of 19 plant species identified using pollen analysis. At present, all of them grow within a small area in the Biya River basin, downstream from Teletskoye Lake (Altai Mountains). Therefore, the current climatic characteristics of this region should be similar to those of the Desna River basin at the time when the studied paleoflora existed (ca. 14–15 14C k.y. B.P.). Method of Estimating Ancient Continental Ice-Sheet Volume Analysis of recent continental and mountain glaciers shows that their shapes can be approximated by ellipses in both the horizontal and vertical planes (Kapitsa, 1958):

x2 y2 + = 1. b2 a 2

(1)

Here, a is the half-length of the short axis of the ellipse in a horizontal plane (in a vertical plane, it corresponds to the maximum height Hm of the glacier), and b is the half-length of a long axis (in a vertical plane, it is the distance from the glacier’s center to its border Lm along a given transect), while x and y are corresponding running coordinates from the ellipse’s center. The unit volume Vu of a glacier’s vertical transect (for the unit width of a glacier) is:

Vu =

π H m Lm ± ε . 4

(2)

Here, ε stands for the volume of initial relief of the glacier’s base, which is neglected in the following equations. An empirical relationship exists between the maximum height Hm (in km) of modern sheet glaciers and the length of a short-axis transect Lm in km (Khodakov, 1982):

H m = ks L m .

(3)

According to the recommendations of Khodakov (1982), the shape coefficient ks in this formula was 0.094 (with the scatter

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from 0.075 to 0.12) for the last Fennoscandian glacier at the “cold” stage of its advance. For the “warm” stage of glacier retreat, he recommends using a coefficient ks equal to 0.061 (0.048–0.077), which yields a flatter glacier. For the stage of glacier retreat, the “dead-ice” glacier model also can be used, with “active” glacier in the center and a marginal zone of melting “inactive” ice; this model follows from geomorphologic considerations (Chebotareva and Faustova, 1982; Faustova, 1984). Therefore, in a short-axis vertical transect for two dated positions of a glacier’s border with a lag Δt (yr) and with distances Lm1 and Lm2 (km), the unit rate of change in “cold” or “warm” glacier ice volume (in km2 yr–1) is:

ΔVu π (ks2 Lm2 − ks1Lm1 ) . = Δt 4Δt 32

32

(4)

For the “dead-ice” glacier model, this unit rate of volume change is:

Methods of River Paleodischarge Estimation The two main ways of calculating paleodischarge from river channel morphology and texture are by either hydraulic or regime equations, both first used by Dury (1965). The advantages and disadvantages of these methods are discussed in Sidorchuk et al. (2008). In the regime equations approach (used here), the relationships between channel morphology and flow hydrology must be known. Our investigation of the rivers on the Russian Plain and in western Siberia (Sidorchuk et al., 2001, 2008) shows that it is important to use a broad range of river sizes and river basin landscapes to work out the empirical formula that suits the purposes of paleohydrological reconstructions. We used ~450 sections of rivers in northern Eurasia with mean annual discharge Qa from 1 to 13,000 m3 s–1 and channel bankfull width Wb from 15 up to 3000 m, and the drainage basins were situated in a variety of landscapes from steppe to tundra. Based on these data, the relationship took the form: Qa = 0.012y0.73Wb1.36.

ΔVu = Δt

π (k L

32 s2 am2

−k L

4Δt

32 s1 am1

)

+ Hd

Ldm2 − Ldm1 . Δt

(5)

Here, Hd is “dead-ice” thickness, and indices “a” and “d” refer to “active” and “dead” (inactive) parts of a glacier. In the case of decrease in glacier volume, this unit rate would be negative and equal to the sum of the positive snow accumulation rate I on the glacier’s surface and the negative rate of meltwater drainage ΔWu/Δt from the glacier plus snowice evaporation rate E from the glacier’s surface. The unit rate of water drainage for a unit width of the glacier (meltwater runoff in km2 yr–1) would therefore be (here ice volumes are recalculated into water volumes with the coefficient 0.9):

ΔWu 0.9ΔVu = + I −E. Δt Δt

(6)

The last two terms in Equation 6 are difficult to estimate for Quaternary glaciers. Unit (for a unit width) snow accumulation rate I (in km2 yr–1) for recent sheet and mountain glaciers (Khodakov, 1982) is related to their size Lm (km):

I = ki L2m3 ,

(7)

⎛ L 2 3 + Lm2 2 3 ⎞ ΔWu 0.9ΔVu = + ki ⎜ m1 ⎟−E. Δt Δt 2 ⎝ ⎠

(8)

and

Coefficient ki for recent glaciers varies in a broad range from 0.001 to 0.005, and therefore its mean value of 0.00224 is recommended for use (Khodakov, 1982).

(9)

Parameter y is inversely related to the seasonal flow variability and represents the ratio between the mean annual discharge Qa and the mean maximum discharge Qmax: y = 100(Qa/Qmax).

(10)

The range of parameter y is from 4 to 5 for rivers with high seasonal flow variability to more than 20 for those with low seasonal flow variability, and up to 100 for rivers with stable flow, such as those draining large lakes. An increase in the flow variability (a decrease in y) generally causes an increase in floodplain height and flow concentration in a single channel with larger bankfull width. Flow variability depends on the basin area F (km2): y = aF 0 .125.

(11)

Parameter a in Equation 11 reflects the geographical distribution of climatic flow variability independent of river basin size. It can be calculated from measured mean annual discharge Qa, mean maximum discharge Qmax, and basin area F. Parameter a typically varies between 1.5 for river basins with high seasonal flow variability to over 4 for river basins with low variability. For paleolandscapes, parameter a for each paleochannel is estimated using recent fluvial analogues. It is then possible to calculate y for paleolandscapes with Equation 11, the mean annual discharge Qa from the paleochannel width with Equation 9, and the mean maximum discharge Qmax with Equation 10. All variables in Equations 9–11 can be obtained from maps and space images, as well as from Hydrological Service measurements. Estimation of coefficient a in Equation 11 for the past requires knowledge of this relationship for the former landscape or its modern hydroclimatic analogue. Geographic influences on river flow bring about similar hydrological regimes for rivers in

Surface runoff to the Black Sea from the East European Plain similar landscapes (Evstigneev, 1990). Geographic controls over river flow and their applications to paleohydrology lead to the principle of paleogeographical analogy (Sidorchuk and Borisova, 2000), which states that the hydrological regime of a paleoriver within a paleolandscape must have been similar to that of a present-day river within the same type of landscape. Therefore, the hydrological regimes of modern rivers in a certain type of landscape can be used to estimate the paleohydrological regime in the same type of paleolandscape. STUDY AREA General Characteristics of the Black Sea Drainage Basin within the East European Plain The Black Sea drainage basin covers the area of ~1,240,000 km2 and occupies ~40% of the East European Plain (Fig. 1). The basin is

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drained by typical lowland rivers: 75% of the territory is situated at the elevations less than 250 m above sea level (asl), and only 2% of the territory is at elevations above 500 m (in the south and southeast, near the Crimean and Caucasian Mountains). The total length of the permanent drainage net of 72,500 rivers in the East European part of the Black Sea basin is 344,258 km (Domanitskiy et al., 1971). The main part of the basin (74.7%) belongs to the Dnieper River (504,000 km2) and the Don River (422,000 km2). All other rivers drain ~23.3% of the basin: the Dniester River (72,000 km2), the Southern Bug River (63,700 km2), the Kuban River (58,000 km2), and other smaller rivers (together 120,000 km2). The large basin size, extending from ~57°N to 44°N, and from 23°E to 41°E, incorporates substantial climate and landscape variability. Mean annual temperature increases from 3 to 9 °C north to south, while annual precipitation decreases from 600 to 300 mm in the same direction. Surface runoff decreases from 200 to 10–20 mm from the north to the south, and then

Figure 1. The East European part of the Black Sea drainage basin. Key: (1) river basin boundaries; (2) region boundaries (see the text for region descriptions); (3) data sites with large paleochannel remnants.

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increases again in the southern mountains. The main part of the basin, including the Dniester River basin, the northern parts of the Dnieper and Don basins, and the Kuban River basin, would be covered with broad-leaved forest under natural conditions. The lower Dnieper and Don basins would be covered by forest-steppe and steppe vegetation under natural conditions. Now, ~70% of the basin is plowed and used for agriculture.

The recent annual flow of all the rivers to the Black Sea from the East European Plain is ~110 km3 (83 km3 from the Dnieper and Don Rivers) or ~30% of the runoff from the total drainage area of the sea. The water regime of the large rivers has been substantially changed by the construction of a system of hydroelectric dams with large reservoirs. Small rivers are often regulated by chains of ponds.

Figure 2. The region with the late Valdai (late Weichselian) glacial and periglacial features. Key: (1) deposits of proglacial lakes; (2) sandy fluvioglacial deposits; (3) meltwater blow-out channels; (4) present-day direction of flow; (5) boundaries between ice sectors. Boundaries of the Last Glaciation stages and keys 1 and 2 are after Faustova and Chebotareva (1969). Key 3 is after Kvasov (1979) with corrections based on space images.

Surface runoff to the Black Sea from the East European Plain

The main geomorphologic features used in hydromorphological reconstruction belong to (1) glacial and periglacial topography (as well as glacial and periglacial deposits), which allows one to trace the boundaries of the former sheet glaciers on the East European Plain, and (2) fluvial topography (with fluvial deposits), which can be used for former runoff estimates. These features, dated to the LGM and LGT, can be found within three regions of the Black Sea basin (see Fig. 1). Region I with the Late Valdai (Late Weichselian) Glacial and Glaciofluvial Features Glacial and periglacial topography and deposits of the Fennoscandian ice sheet in the Dnieper River basin were investigated mainly in 1960s and 1970s, when several large monographs were published (Gerasimov, 1969; Chebotareva and Makarycheva, 1974). The position of the southern boundary of the Fennoscandian ice sheet here during the LGM has been generally confirmed by recent works (Velichko et al., 2004; Svendsen et al., 2004). It is firmly established that during its maximum extent, the last ice sheet covered only the northernmost part of the Upper Dnieper River basin (region I in Figs. 1 and 2). Glacial topography is represented by a system of moraine hills and ridges often clearly bordering the former ice lobes. Three main bands of such moraine morphology, dated with 14C and pollen analysis of under-moraine deposits, show the position of the glacier boundary (Fig. 2) during the Bologoye/Brandenburg stage (ca. 18 14C k.y. B.P.), the Edrovo/Frankfurt stage (ca. 17 14C k.y. B.P.), and the Vepsovo/Pomeranian stage (ca. 15.5 14C k.y. B.P.)—all the dates are after Chebotareva and Makarycheva (1974), the last being entirely beyond the Black Sea basin. Taking the center of the southern half-ellipse of the Fennoscandian glacier at the center of the northern part of the Gulf of Bothnia, the distance (Lm) to the glacier border (at the headwaters of the Berezina River) was ~1230 km at the Bologoye stage, ~1180 km at the Edrovo stage, and ~1100 km at the Vepsovo stage (Table 1). Each stage was characterized by an “active” phase of the glacier advance and by an “inactive” phase of the glacier retreat, when a marginal zone of “dead-ice” could exist. Another important morphological and depositional feature is represented by the fluvioglacial plains formed by fluvioglacial streams, terraces, and deposits of ice-dam lakes and river terraces related to them. These features are also dated with 14C and by their correspondence to the glacial topography. According to Faustova and Chebotareva (1969) and Chebotareva and Makarycheva (1974), the “glacial” part of the Black Sea drainage basin achieved its maximum area during the Bologoye (Brandenburg) stage ca. 18 14C k.y. B.P. At that time, a chain of ice-dam lakes (Kvasov 1979), or a series of short-lived glacial lakes (Mangerud et al., 2004), formed a broad band between the glacier front and the main pre–last glaciation water divide of the East European Plain. The chain of lakes stretched from the headwaters of the Dnieper River in the east to the upper part of the Neman River

TABLE 1. MELTWATER SUPPLY TO THE DNIEPER RIVER HEADWATERS DURING THE EARLY STAGES OF THE FENNOSCANDIAN ICE-SHEET RETREAT F Hm E Sector of drainage Ice-sheet stage Lm I tcal Lf 0.9 ΔVu/Δt Δt ΔWu/Δt 3 1 3 1 3 1 (km yr– ) (km yr– ) (km3 yr–1) (km) (km) from ice sheet (yr B.P.) (km) (km) (km yr– ) (yr) “Cold” glacier scenario Bologoye 1230 670 410,000 3.3 21,500 86 25 Edrovo 1180 670 390,000 3.2 20,300 83 24 A+B+C 1200 54 115 Edrovo 1180 440 260,000 3.2 20,300 55 15 B+C 1800 26 65 Vepsovo 1100 440 245,000 3.1 18,500 53 15 Edrovo 1180 270 160,000 3.2 20,300 34 10 Vepsovo 1100 270 150,000 3.1 18,500 33 9 C 1800 16 40 “Warm” glacier scenario Bologoye 1230 670 410,000 2.1 21,500 86 25 A+B+C 1200 35 96 Edrovo 1180 670 390,000 2.1 20,300 83 24 Edrovo 1180 440 260,000 2.1 20,300 55 15 Vepsovo 1100 440 245,000 2.0 18,500 53 15 B+C 1800 17 55 Edrovo 1180 270 160,000 2.1 20,300 34 10 Vepsovo 1100 270 150,000 2.0 18,500 33 9 C 1800 10 34 “Dead-ice” glacier scenario Bologoye 1230 670 410,000 0.5 21,500 86 25 A+B+C Edrovo 1180 670 390,000 0.5 20,300 1200 83 24 7 68 Edrovo 1180 440 260,000 0.5 20,300 55 15 B+C 1800 3 42 Vepsovo 1100 440 245,000 0.5 18,500 53 15 Edrovo 1180 270 160,000 0.5 20,300 34 10 Vepsovo 1100 270 150,000 0.5 18,500 33 9 C 1800 2 26 Note: See explanations of the indexes in the text. Note that all unit rates are multiplied on the sector half-width.

Main Geomorphologic Regions in the Black Sea Basin

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basin in the west (Fig. 2). The lakes were connected by channels and formed a pool-step system with a general slope from east to west. Mangerud et al. (2004) did not assume any southward meltwater drainage from these lakes. Nevertheless, there is clear geomorphic evidence of meltwater blow-outs through a number of river valleys. During the Bologoye stage, when this lake system was separated from the marginal valleys in the Polish and German lowlands, the meltwaters drained to the south through the lowest parts of the main water divide. At different times, the flows used different passes, so that a complicated pattern of blow-out valleys was formed at the headwaters of the Dnieper, Dvina, and Neman Rivers. Kalicki (1995) even suggested a new “Dnieper-type” of paleochannel with multiple old valleys. On the whole, there were three main routes of meltwater drainage: through upper Dnieper valley, through the upper Berezina River, and through the Pripyat River, each corresponding to one of three sectors of the ice sheet (A, B, C in Fig. 2). At that time (ca. 17 14C k.y. B.P.), the first terrace in the Upper Dnieper basin was formed (Kalicki and San’ko, 1992). This terrace is ~1800 m wide in the Upper Dnieper valley, ~1200 m wide in the Upper Berezina valley, and ~1600 m wide in the valley of the Shara River—a tributary of the Neman River that connected a paleolake at the Upper Neman basin with the Pripyat River valley (Fig. 2). The sizes of large meanders in these paleochannels correspond well to their widths. With the last Fennoscandian glacial retreat from the boundaries of the maximum (Bologoye) stage, the area of the additional “glacial” part of the Black Sea basin (F) decreased from 410,000 km2 (sectors A + B + C) to 260,000 km2 (sectors B + C), as the width of the ice-sheet front (Lf) was reduced from 670 to 440 km (see Table 1). It happened when the paleo–Upper Neman River became connected with the marginal valleys in the Polish and German lowlands through the Neman paleovalley, and meltwater flow into the Pripyat River stopped. The meltwater fluvioglacial streams and ice-dam lakes continued to drain into the Upper Dnieper and Upper Berezina only in the eastern part of the system. During the Edrovo stage, the area of the additional “glacial” part decreased again, from 260,000 km2 to 150,000 km2 (sector C), when meltwater drainage into the Upper Berezina River stopped. After the Dvina River valley formed during the Vepsovo stage, the entire system of meltwater drainage shifted into the Baltic Lake and into the western marginal valleys, so that the Black Sea lost its connection with the meltwater source. Region II with the Late Glacial Large Alluvial Paleochannels Well-preserved fragments of large meandering paleochannels can be distinguished on large-scale maps and space images. In our former investigations, 16 such fragments over rivers >200 km long were described in the Black Sea basin (see table 1 in Sidorchuk et al., 2001). The use of Landsat-7 images with 15 m resolution allowed us to find an additional 35 fragments within river valleys with basins greater than 5000 km2 (Table 2 in this paper). Although a future increase in image resolution will potentially increase the number of large paleochannels traced, currently available data on the Black Sea drainage basin

provide the possibility of analyzing the general distribution of such features. Remnants of the large alluvial paleochannels are wellpreserved in the basins of the tributaries of the Dnieper and Don rivers (region II in Fig. 1). These remnants are mainly situated at the level of the modern floodplain. In the north, the necks of paleomeanders form fragments of the first terrace, while the point bars and filled paleochannels constitute the floodplains of the recent valleys. In several river valleys, the paleochannel fragments are partly situated at the low terraces and partly included within the recent floodplain. In the southern part of the region, all elements of paleochannels are included within the recent floodplains. Therefore, rivers in this region are characterized by very high ratios of floodplain to channel widths. The relationships between river channel plan geometry and flow discharge are of prime importance for paleohydrological reconstruction. The large paleochannels (in terms of their width and meander wavelength) were up to 15 times larger than the recent channels in the same river basin. Such sizes indicate large surface runoff at the time of paleochannel formation. Two key sites were investigated in this region: one in the Dnieper River basin (Seim and Svapa Rivers), and the other in the Don River basin (the Khoper River). Paleochannels in the Seim and Svapa River Valleys (Dnieper River Basin) The floodplain of the Seim River and its main tributaries is characterized by a sequence of arcs with their radii of curvature exceeding that of the modern river channel bends by an order of amplitude (Figs. 3 and 4). Systems of natural levees and large abandoned oxbows are well defined on aerial photographs. Systems of levees and hollows with relative relief of 0.5–1.5 m reflect steeply curved meandering paleochannels with wavelengths of 3.3–6.5 km and widths of 350–700 m (for the recent Seim River, these are 0.2–1.0 km and 20–100 m, respectively). The floodplains with remains of large paleochannels lie largely at 2.0– 2.5 m above low water level (LWL) and usually become inundated during floods. Only the tops of the highest levees are as high as 3–3.5 m and usually remain above flood level. The investigated fragment of the large paleochannel of the Seim River near Kudintsevo village is the highly curved meander with a wavelength of 6 km. The texture of infilling of this large paleochannel was investigated by coring along the profile A′–A″ in the upper part of the bend (Fig. 3). The paleochannel cross section has an asymmetrical triangular shape with the deepest part (~10 m below LWL) near the steep concave bank of the paleomeander. The paleochannel fill includes three main units. The lowermost unit is fine silty sand 5–7 m thick, which belongs to the initial stage of the paleochannel infill. This sand is overlain by gray clay and silt accumulated largely in the oxbow lake on the floodplain. Deposits at the base of this unit were radiocarbon dated to 12,630 ± 70 and 13,800 ± 85 yr B.P., based on bulk organic matter (samples Ki-6985 and Ki-6984). We assume that the latter date indicates the time shortly after the macromeander

Surface runoff to the Black Sea from the East European Plain TABLE 2. PALEORUNOFF RECONSTRUCTIONS BASED ON THE SIZES OF LARGE 14 ALLUVIAL PALEOCHANNELS IN THE BLACK SEA DRAINAGE BASIN CA. 14–15 C k.y. B.P. W Q Latitude Longitude F λ River 2 3 (m) (m /s) (°N) (°E) (km ) (m) Aidar 49.01 3 8 .9 6 7015 845 1 74 48 Berezina 54.02 28.87 9557 991 260 74 Bitiug 51.03 40.06 7695 1177 211 69 Buzuluk (Khoper) 50.64 42.79 6390 1206 182 62 Desna 51.13 31.03 84,183 2145 385 194 Desna 51.44 31.95 71,145 1806 278 136 Desna 52.45 33.56 24,989 1513 297 115 Dnieper 54.97 32.98 7340 1997 136 Egorlyk 45.99 41.30 1 1 ,2 2 0 938 50 Egorlyk 45.87 41.40 8 707 845 171 48 Iput' 52.47 31.39 9400 809 41 Kalitva 48.42 40 . 96 10 , 14 0 80 5 133 40 Khoper 51.27 42.33 19,329 1906 328 140 Khoper 52.46 43.67 8981 1358 184 71 50.19 43.73 29,612 1810 213 107 Medveditsa (Don) Medveditsa (Don) 51.44 44.86 7582 973 184 55 Nerussa 52.44 33.90 5346 890 209 56 Orel' 48.97 34 . 4 9 91 58 1 26 5 187 68 Orel' 48.93 34 . 5 9 91 58 9 73 203 60 Orel' 49.16 34 . 9 3 74 76 9 12 152 46 Orel' 49.13 35 . 1 2 56 53 7 92 153 41 Psel 49.23 33.68 22,1 58 12 03 152 62 Psel 49.74 33.78 14,6 92 841 162 49 Psel 50.18 33.97 11,7 35 712 122 35 Psel 50.56 34.44 94 17 67 1 1 21 33 Psel 51.04 35.22 65 87 67 0 30 Ros' 49.49 31.48 11,021 3701 458 267 Sal 47.33 41.34 20,523 1392 288 105 1936 286 130 Samara (Dnieper) 48.64 35.37 19,903 Seim 51.39 33.41 27,070 1938 234 118 Seim 51.28 33.88 2 2,257 1602 220 97 Seim 51.48 34.78 19,678 1673 312 124 Seim 51.69 35.23 11,418 1924 349 140 18,706 742 152 44 Severskiy Donets 49.33 36.84 Snoc 51.66 31.64 8281 2311 167 Sozh 52.35 30.95 39,336 1693 380 154 Sozh 53.86 31.80 6630 225 77 Styr' 50.77 25.34 7201 1762 114 Sula 49.65 32.71 1 8 , 00 9 1 5 95 26 3 106 Sula 49.62 32.88 1 8 , 00 9 1 6 62 27 2 112 Sula 49.81 32.90 1 5 , 29 5 1 7 27 19 0 91 Sula 50.22 33.32 60 9 8 1 26 0 238 77 Tersa 50.87 44.11 6589 963 171 51 Udaj 50.21 3 2 .7 6 65 79 1 2 04 2 10 69 Volchia 48.11 3 6. 0 7 9627 77 1 13 9 40 Vorona 51.78 4 2.37 12,746 1 434 252 93 Vorskla 49.00 34 .16 14,433 12 93 256 89 Vorskla 49.04 34 .33 14,433 12 16 238 81 Vorskla 49.49 34 .58 11,124 83 8 1 15 38 Vorskla 49.73 34 .63 9 129 1259 207 72 Vorskla 50.14 34 .74 6 220 75 9 133 37 Note: F—catchment area, λ—mean paleomeander step (half-wavelength), W—mean paleochannel width, Q—mean annual discharge, X—mean annual runoff depth.

was abandoned. By the beginning of the Holocene, the oxbow lake had been transformed into a fen, and mineral deposition was followed by peat accumulation. The thickness of the peat layer is up to 2 m. In the Svapa River valley, the paleochannel formed steeply curved meanders with a mean wavelength of 2.8 km and a mean width of 300 m (for the recent Svapa River, these are 0.14–

9

X (mm) 216 245 282 307 72 60 145 583 142 175 136 124 228 248 114 229 328 233 207 194 229 89 104 94 110 146 763 161 205 138 137 199 387 74 637 124 364 500 186 195 188 396 245 329 130 230 194 177 107 249 185

0.6 km and 15–60 m, respectively). The paleochannel fragment near Semenovka village is clearly expressed in the modern topography (Fig. 4). Its surface lies only 4 m above the modern LWL, so that it is submerged during high floods. Coring reveals that the channel trough assumes a box-shaped profile. The top of the lower layer of fine- and medium-grained channel alluvial sands lies 1.5–2.5 m below the modern LWL. Silt and clay deposits

10

Sidorchuk et al.

with lenses of clayey sand fill the trough. Accumulation of finegrained sediments began because of the abandonment of the paleochannel in the Oldest Dryas (14,030 ± 70 yr B.P., Ki-6997), and continued during the Bølling and the Allerød (12,360 ± 110 yr B.P., Ki-6999 and 11,755 ± 80 yr B.P., Ki-6996). At the end of the Late Glacial to beginning of the Holocene, the paleochannel almost entirely dried up, and the rate of deposition became very low. Peat formation started in the late Preboreal (9120 ± 70 yr B.P., Ki-6995; 9300 ± 120 yr B.P., GIN-11951). Paleochannels in the Khoper River Valley (Don River Basin) One of the best preserved systems of large paleochannels with bankfull channel width of 0.8–1.4 km, maximum depth of 9 m, and mean wavelength L of 5 km, is found on the ancient floodplain of the Khoper River near Povorino (Fig. 5). At present, the Khoper River near Povorino has a channel width of 60 m, maximum depth of 4 m, and a meander length up to 360 m. Radiothermoluminescence dating (Vlasov and Kulikov, 1988) of the bottom deposits (17 ± 4 ka, RTL-808) shows that the large paleochannels were formed during the Late Glacial. The

thickness of their subsequent infilling varies at different parts of the paleochannel. The system of channels that follows the right bank of the valley has been rarely flooded after it was abandoned, so that the former bottom of the paleochannel is locally exposed. The maximum thickness of the channel-filling deposits there does not exceed 1.5–3.0 m (Fig. 5). There are large eolian dunes on the terrace at the macromeander neck. The system of channels along the left bank of the valley is situated nearer to the present river. The pools of the large paleochannel with maximum depth of 9–11 m were completely filled in by fine-grained alluvium, beginning from 11.3 to 10.8 14C k.y. B.P. A system of smaller paleochannels (although still larger than the recent Khoper River channel) was preserved at the outspread of the floodplain in the valley bend (see Fig. 5, cross-section C–C′). The channel width was ~200 m, and the meander wavelength was ~1200 m. These channels were active before Boreal time, presumably during the Younger Dryas. Deposition of floodplain sediments during the Holocene was concentrated in a narrow belt, 1 km wide, along the river channel, where sandy natural levees up to 3.5 m above LWL were formed.

Figure 3. Geological section and space image across the large meandering channel of the paleo–Seim River near Kudintsevo village.

Surface runoff to the Black Sea from the East European Plain Region III with Large Incised Paleochannels Region III (see Fig. 1) occupies the high western part of the East European Plain (Volyno-Podolsk Upland), the near–Black Sea plain, and pre-Caucasus highlands. Rivers are often incised here, presumably because of tectonic uplift. The channels of these rivers form large bends with systems of Quaternary terraces at their necks. The ages and origins of these large bends are mostly unknown, and their paleohydrological signal is unclear. These features are not used in further runoff estimates. RESULTS OF THE ESTIMATES OF PALEOGEOGRAPHIC CONDITIONS Hydroclimatic Parameters of River Development since the Late Pleniglacial To reconstruct the main climatic indexes in the Black Sea basin, composition of fossil floras derived from palynological data (the method of “arealograms”) was analyzed. Fossil floras used for paleoclimatic reconstructions were derived from five sites

11

(Fig. 6). At three of them, palynological studies were conducted on fluvial sediments that filled segments of large paleochannels found on the floodplains of the Seim, Svapa, and Khoper Rivers (Sidorchuk and Borisova, 2000; Borisova et al., 2006). Another of the localities is the Early Man site of Yudinovo, where loamy sediments containing a Late Paleolithic cultural layer were subjected to detailed pollen analyses (Borisova and Novenko, 1999). One more fossil flora includes both pollen and plant macrofossils identified in the so-called “Usvyacha” deposits exposed in several outcrops within the Dvina River valley near the villages of Sloboda and Drichaluki (Velichkevich, 1982; San’ko, 1987). In the course of pollen analysis, an attempt was made to achieve the highest possible taxonomic resolution. Pollen identifications have been made to species or genus levels for arboreal plants as well as certain groups of herbaceous plants. When identifications were possible only to the family level, the finds in question were not included in the resulting lists of fossil floras (Table 3). On the whole, these paleobotanical data proved to be sufficient to locate modern geographical analogues to 12 fossil floras and, therefore, to estimate climatic changes that occurred within the

Figure 4. Geological section and space image across the large meandering channel of the paleo–Svapa River near Semenovka village.

Figure 5. Geological sections and space image across the large meandering channels of the paleo–Khoper River near Povorino. RTL—radiothermoluminescence.

12 Sidorchuk et al.

Surface runoff to the Black Sea from the East European Plain

13

Figure 6. Location of the modern region analogues to the Last Glacial Maximum and Late Glacial time fossil floras. Key: (1) key sites (Sm— Seim, Sv—Svapa, Kh—Khoper, Yu—Yudinovo, Sl—Sloboda/Drichaluki); (2) modern region analogues of the fossil floras (number of a fossil flora and its 14C age in k.y. B.P.); (3) corresponding hydrological region analogues in the lowland areas (A—Bol’shezemel’skaya Tundra, B— Lena Plateau).

East European part of the Black Sea drainage basin since ca. 18 14 C k.y. B.P. The accuracy of the match between a modern geographical analogue and a fossil flora depends not only on the richness of the latter, but also on the knowledge of the modern geographical distribution of each plant genus or species. In some cases, the area of a region analogue remains rather large and includes different types of vegetation associations. Therefore, in this study, all such regions were located on the map (see Fig. 6) and additionally checked afterward against the types of plant communities, reconstructed from the pollen assemblage of a given sample. Sometimes, such comparisons enabled us to reduce the area of the region analogue, but even then, the resulting areas remained large enough to show a substantial variability of hydroclimatic characteristics within them. The reconstructed ranges of climatic parameters are shown in Figure 7 as shaded boxes, their vertical size corresponding to the time intervals characterized by each fossil flora. The earliest fossil flora (flora 1 in Table 3 and in Figs. 6 and 7)—derived from the palynological and plant macrofossil studies of the Sloboda and Drichaluki sections—belongs to the relatively cold stage of the Late Pleniglacial, dated by radiocarbon to ca. 17–18 ka (Velichkevich, 1982; San’ko, 1987). The flora includes several of the typical Arctic and Arctic-mountain species,

which at present grow in various European and west Siberian tundra and forest-tundra communities, along with some xerophytes tolerant to low winter temperatures, such as Ephedra distachya. It also includes trees and shrubs growing presently in regions with cold and highly continental climate, mainly in Siberia (Larix sp., Alnaster fruticosus). Such a complexity of flora is highly typical of the glacial floras in northern Eurasia (Grichuk, 1969). The region currently inhabited by the species of fossil flora 1 lies south of the East Sayan Mountains, in the upper part of the Oka River basin. The area has a cold climate with mean January air temperature of –21 to –22 °C and mean July air temperature of 8– 10 °C, which are characteristic of mountain tundra in this region. Because of permafrost, the runoff coefficient should be very high (more than 0.8). The mean annual precipitation is 400–600 mm, and the calculated runoff depth is 350–500 mm (Table 4). The Early Man site of Yudinovo is located on the first terrace of the Sudost’ River, a tributary of the Desna River in the Dnieper basin. Both the geomorphological position of the site and a series of radiocarbon dates based on the materials from the cultural layer indicate that the site was inhabited approximately from 14.5 to 14 ka (Svezhentsev, 1993). According to palynological data, fossil flora of this interval consists mainly of forest and meadow plants. It includes Boreal and Arctic-Boreal trees and shrubs, forest club-mosses (Lycopodium clavatum, L. selago, and

14

Sidorchuk et al. TABLE 3. THE LATE GLACIAL AND THE HOLOCENE PALEOFLORAS FOUND IN THE BLACK SEA DRAINAGE BASIN

Site names:

Sloboda/ Drichaluki Fl 1

Yudinovo

Seim

Seim

Numbers of fossil floras: Fl 2 Fl 3 Fl 4 Radiocarbon ages of the (17–18) (14–14.5) 13.8 12.6 fossil floras (k.y. B.P.)#: I. Trees and shrubs (a) Micro- and mesothermal species of the continental regions Abies sibirica * Alnaster fruticosus *† * * * § Larix sp. ** * Pinus sibirica * * (b) Species with broad ecological tolerances Betula alba * * * * Betula humilis * * * Picea abies * * * * Pinus sylvestris * * * * (c) Relatively thermophile mesophytes Acer tataricum Alnus glutinosa A. incana * * * Corylus avellana Quercus robur Tilia cordata Ulmus sp. U. campestris Viburnum opulus II. Arctic-Alpine microthermal mesophyte species Betula nana ** * Dryas octopetala ** Gastrolychnis apetala ** Lycopodium pungens * Polygonum viviparum ** Potentilla cf. nivea ** Salix herbacea ** Salix cf. polaris ** Selaginella selaginoides ** * III. Herbaceous plants associated with various forest communities Botrychium ramosum * Humulus lupulus Lycopodium annotinum L. clavatum * * * L. complanatum * L. selago * L. tristachyum * Thalictrum minus T. simplex * Pteridium aquilinum * IV. Xerohalophytes Atriplex cana * A. pedunculata A. verrucifera Chenopodium glaucum * C. acuminatum C. chenopodioides Kochia prostrata * * Plantago cornuti Salsola sp. * S. soda V. Xerophytes Helianthenum sp. * * * Eurotia ceratoides Ephedra (non-distachya) * * * * Ephedra distachya Kochia scoparia * VI. Psammophytes and plants growing on eroded soil Amaranthus sp. Atriplex tatarica Centaurea cyanus Chenopodium album C. botrys Linaria sp. Spergula sp.

Svapa

Seim

Khoper

Seim

Svapa

Svapa

Svapa

Svapa

Fl 5

Fl 6

Fl 7

Fl 8

Fl 9

Fl 10

Fl 11

Fl 12

12.4

12.2

11.9

11.5

(10.5)

9.8

7.5

4.9

*

* * * *

* *

* * * * * * * *

*

* * * *

* * * *

* *

* * * * *

* *

* * *

* * * *

*

* *

*

*

*

* * * *

* * * * * * * * *

*

*

* * *

*

* * *

* * *

* * *

* * *

* * * * *

* * *

*

*

* *

* *

* *

* * * * *

* * *

* * * *

* * *

* *

* (continued)

Surface runoff to the Black Sea from the East European Plain

15

TABLE 3. THE LATE GLACIAL AND THE HOLOCENE PALEOFLORAS FOUND IN THE BLACK SEA DRAINAGE BASIN (continued) Site names:

Sloboda/ Drichaluki Fl 1

Yudinovo

Seim

Seim

Numbers of fossil floras: Fl 2 Fl 3 Fl 4 Radiocarbon ages of the (17–18) (14–14.5) 13.8 12.6 fossil floras (k.y. B.P.)#: VII. Steppe and meadow plants Cannabis sp. Fagopyrum sp. * Plantago ramosa Rumex acetosella Botrychium lanceolatum * B. simplex Bupleurum sp. Papaver nudicaule ** Plantago lanceolata * Sanguisorba officinalis Scabiosa sp. * Thalictrum foetidum Valeriana sp. * VIII. Plants of wet meadows, water margins, and mires (helophytes) Alisma gramineum A. plantago-aquatica * Calystegia sp. Equisetum palustre E. scirpoides E. variegatum Filipendula ulmaria Menyanthes trifoliata Myosoton aquaticum ** Polygonum amphybium Ranunculus reptans ** Sparganium sp. * * S. hyperboreum ** Sagittaria sagittifolia Thalictrum angustifolium T. flavum Typha angustifolia T. latifolia Urtica sp. IX. Aquatic plants Batrachium sp. ** Lemna sp. Myriophyllum sp. * M. spicatum * M. verticillatum Nymphaea alba Nymphaea candida Potamogeton filiformis ** P. natans P. perfoliatus ** P. vaginatus ** # Age estimations based on interpolation are shown in parentheses. † Plants identified by their pollen or spores. § Plants identified by macrofossils.

others), Pteridium aquilinum, and other mesophile plants (see Table 3). Such a floral composition suggests relatively humid conditions during the interval in question. A region analogue for fossil flora 2 lies in the Altai Mountains, in the middle reaches of the Biya River, on the east-facing slopes (Fig. 6). In this region, pine and birch forests and spruce–fir–Siberian pine mountain taiga forests occur along with wet meadows. The area is characterized by milder and wetter climatic conditions compared to the region analogue 1: the mean January air temperature there is –16 °C, and the mean July temperature is 16–17 °C. The mean annual precipitation is 700–800 mm, and the runoff depth is 500– 550 mm (see Table 4).

Svapa

Seim

Khoper

Seim

Svapa

Svapa

Svapa

Svapa

Fl 5

Fl 6

Fl 7

Fl 8

Fl 9

Fl 10

Fl 11

Fl 12

12.4

12.2

11.9

11.5

(10.5)

9.8

7.5

4.9

* * *

*

* * * * *

*

*

*

*

* *

*

* *

* * *

*

* * *

* *

*

* * *

*

* *

*

*

*

*

*

* *

*

*

*

* * *

* *

* *

* *

* *

* * *

* * *

* * * *

Sediments of the initial stage of filling of the large paleochannel of the Seim River (core S-4 in Fig. 3; 13,800 ± 85 yr B.P., Ki-6984) generally correspond to the beginning of the Oldest Dryas. Fossil flora 3, derived from these sediments, combines cryophile (Alnaster fruticosus, Selaginella selaginoides) and xerophile (Ephedra sp.) plants, inhabitants of the boreal forest, steppe, meadow, and riverine communities. The closest modern floristic analogue for this assemblage can be found in the forest steppe in the middle reaches of the Irkut River basin, west of Lake Baikal (see Fig. 6). Within this small area, larch and pine forest grow next to southern Siberian meadow steppes, with patches of spruce forest in the river valleys. The area is characterized by a

Figure 7. Estimations of the main climatic indexes (deviations from modern values) based on the composition of fossil floras (numbers of the floras are as in Tables 3 and 4 and Fig. 6). PB—Preboreal; DR3—Younger Dryas; AL—Allerød; BØ—Bølling; DR1—Oldest Dryas.

16 Sidorchuk et al.

2.0 3.5 3.0 2.6 3.4 1.6 1.6 1.8 2.0 1.2 1.4 0.8 350 to 400 225 to 275 175 to 225 240 to 290 25 to 125 40 to 90 75 to 125 40 to 190 –10 to 45 20 to 70 –50 to 15

140 to 290 350 to 500

500 to 550 350 to 400 300 to 350 350 to 400 150 to 250 150 to 200 200 to 250 150 to 300 100 to 155 130 to 180 65 to 125 125 to 225 –125 to –75 –175 to –125 –125 to –75 50 to 150 40 to 140 125 to 175 –150 to –50 –50 to 50 100 to 150 75 to 125

–200 to 0 400 to 600

700 to 800 425 to 475 375 to 425 425 to 475 600 to 700 500 to 600 675 to 725 400 to 500 500 to 600 650 to 700 625 to 675 –1.5 to –2.5 –3.5 to –2.5 –2 to 0 –4 to 0 –3 to –1 –3 to –4 –0.5 to –1.5 –3.5 to –4.5 –0.5 to 0.5 0 to –1 0 to –1

–7 to –9 8 to 10

16 to 17 15.5 to 16.5 17 to 19 15 to 19 16 to 18 16 to 17 17.5 to 18.5 14.5 to 15.5 18.5 to 19.5 18 to 19 18 to 19 –5.5 to –9.5 –14 to –18 –15 to –19 –13 to –15 –7.5 to –8.5 –6 to –8 –7.5 to –8.5 –9 to –11 –7 to –8 –7 to –8 1 to 2

–13 to –14 –21 to –22

–14 to –18 –22 to –26 –23 to –27 –21 to –25 –15.5 to –16.5 –16 to –18 –15.5 to –16.5 –17 to –19 –15 to –16 –15 to –16 –6 to –7

(17–18)

(14–14.5) 13.8 12.6 12.4 12.2 11.9 11.5 (10.5) 9.8 7.5 4.9

1

2 3 4 5 6 7 8 9 10 11 12

Sloboda/ Drichaluki Yudinovo Seim Seim Svapa Seim Khoper Seim Svapa Svapa Svapa Svapa

ΔRunoff depth (mm/yr) Runoff depth (mm/yr) ΔPrecipitation (mm/yr) Precipitation (mm/yr) ΔTemp. July (°C) Temp. July (°C) ΔTemp. January (°C) Temp. January (°C) No. of fossil flora

Site names Radiocarbon ages of the fossil floras (k.y. B.P.)

TABLE 4. MAIN CLIMATIC INDEXES IN THE BLACK SEA DRAINAGE BASIN IN THE LATE GLACIAL AND THE HOLOCENE

Runoff depth ratio past/recent

Surface runoff to the Black Sea from the East European Plain

17

cold semiarid and extremely continental climate. The mean January temperature is –24 °C, and that of July is ~16 °C. The mean annual precipitation varies between 425 and 475 mm. The region is situated near the boundary of the permafrost zone with a high annual runoff coefficient. The mean annual surface runoff depth is 350–400 mm (see Table 4). Another sample of fluvial deposit fill from the large paleochannel of the Seim River was obtained from the basal part of core S-7 taken in the same profile as core S-4 (Fig. 3). According to the radiocarbon date (12,630 ± 70 yr B.P., Ki-6985), it corresponds to the final part of the Oldest Dryas. The flora of this sample (flora 4 in Table 3) is distinctive for the diversity of its xerophytes and xerohalophytes, which suggest a rapid drying of the topsoil in the watershed areas during relatively warm summers. Similar conditions occur at the present time in southern Siberia. A region analogue for fossil flora 4 lies at the headwaters of the Yenisei River, within an intermountain depression downstream from the confluence of the Biy-Khem and Ka-Khem rivers (Fig. 6). In this area, southern Siberian dry grassy steppes are found next to a mountain forest of Pinus sibirica and Larix sibirica. The area is characterized by extremely cold winters with mean January air temperature from –23 °C to –27 °C, while the summer is warm with mean July temperature being ~18 °C. The annual magnitude of air temperature changes is ~43 °C, and that is ~15 °C greater than at the study site at the present time. The mean annual precipitation is ~400 mm, and calculated runoff depth is ~325 mm. Fluvial deposits filling the paleochannel of the Svapa River were cored near Semenovka village (see Fig. 4). According to the 14 C date (12,360 ± 110 yr B.P., Ki-6999), fossil flora 5, deriving from core SV-1–8, belongs to the beginning of the Bølling interstadial. It includes species of dark coniferous taiga forest (Picea abies, Abies sibirica, Pinus sibirica), and light coniferous and mixed boreal forest (Pinus sylvestris, Pteridium aquilinum, Betula alba), along with species of riverine shrub associations, meadows, psammophytes (e.g., Spergula), and xerophytes. A region analogue for fossil flora 5 lies at the headwaters of the Yenisei River. It is shifted to the north with respect to region analogue 4 (see Fig. 6). In this area, southern Siberian dry grassy steppes are also the main vegetation type, though the role of the dark coniferous mountain forest of Picea, Pinus sibirica, and Abies sibirica is slightly greater here compared to region analogue 4. The area is characterized by milder climatic conditions, with mean January air temperature of –23 °C and mean July air temperature of ~17 °C, which are similar to the reconstruction based on fossil flora 3. Therefore, the runoff coefficient could be the same as at the region analogue of flora 3. The mean annual precipitation is ~450 mm, and the calculated runoff depth is ~375 mm. Palynological data for the fluvial deposits infilling the paleochannel of the second generation in the Seim River valley were obtained from borehole S-11 (see Fig. 3). Flora 6, derived from a sample dated to 12,250 ± 70 yr B.P. (Ki-6987), includes broad-leaved temperate tree species (Quercus, Ulmus, and Tilia cordata), as well as Alnus glutinosa, tree alder, growing on the

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waterlogged ground. Its composition indicates that in the late Bølling, the area was covered by a complex vegetation of forest steppe type, with birch and pine copses and minor participation by broad-leaved trees. The presence of broad-leaved trees in the region implies a complete degradation of permafrost at the end of the Bølling. This is confirmed by the generally mesophile character of the flora and by the presence of pollen of relatively heatdemanding aquatic plants, such as Nymphaea alba (see Table 3). A diversity of the hygro- and hydrophytes and an absence of xerohalophytes in this flora suggest an increase in rainfall. A region analogue for fossil flora 6 lies at the headwaters of the Ufa River in the Southern Ural Mountains (Fig. 6). In this region, meadow steppes come into close contact with the southern Urals pine and birch forests and spruce-fir subtaiga forests, with a minor presence of broad-leaved trees. The area is characterized by significantly milder climatic conditions compared to the region analogue of the early Bølling flora 5: the mean January air temperature there is –16 °C, and the mean July temperature is ~17 °C. The mean annual precipitation is 650 mm, and the runoff depth is ~200 mm, with 70% of the flow passing during the spring flood. The annual runoff coefficient is 0.31, and that for the flood period is ~0.6. The alluvium deposits infilling the meandering Khoper paleochannel were studied near the town of Povorino (see Fig. 5). At the base of core B, a radiocarbon date was obtained for the layer enriched with organic matter: 11,900 ± 120 yr B.P. (Ki-5305). The flora of this horizon includes some arboreal species with broad ecological tolerances, such as Scots pine (Pinus sylvestris) and tree and shrub birch (Betula alba, Betula humilis), but it also includes Siberian pine (P. sibirica), growing in the regions with continental climate, xerohalophytes, and xerophytes (Atriplex pedunculata, Atriplex verrucifera, and others; see Table 3). A region analogue of fossil flora 7 is located in northeastern Kazakhstan, in the Bukhtarma River basin (Irtysh River basin) at the boundary between dry steppe and semidesert and close to the region where communities of shrub birch and open dark coniferous forest are spread on the western slopes of the Altai Mountains (Fig. 6). The mean January air temperature in this region is –17 °C; mean July temperature is ~16.5 °C. The mean annual precipitation varies between 500 and 600 mm. The annual runoff depth is 150–200 mm. The region analogue is situated beyond the permafrost zone but close to its boundary. Fossil flora 8 of the dated sample from section S-11 (see Fig. 3; 11,450 ± 60 yr B.P., Ki-6986) includes many helophytes—plants growing on wet meadows, near the water’s margin, and in shallow water (Sparganium spp., Menyanthes trifoliata, Sagittaria sagittifolia, and others; see Table 3). Composition of this flora indicates that the Allerød warm interval was favorable for the expansion of dark coniferous trees (Picea abies, Pinus sibirica, and Abies) in the Seim River basin. Forest communities were dominated by birch and Scots pine, and larch participated in the pine forest at this period. Of the relatively thermophile trees and shrubs, flora of the Allerød included Ulmus, Tilia, Corylus, and Viburnum. A modern analogue for fossil flora 8 is located at the headwaters

of the Belaya River near the Southern Urals (Fig. 6). The region is currently occupied by forest steppe, where meadow steppes are associated with birch and aspen woods and grow next to (1) larch-pine open forests with steppe elements in the herbaceous cover, and (2) broad-leaved (oak-linden) forests of Southern Ural type. This area is characterized by the mildest climatic conditions for the entire Late Glacial, with mean January air temperature at –16 °C and July air temperature at 18 °C. Mean annual precipitation there reaches 700 mm, and runoff depth is ~225 mm, with 60% of the flow passing during the spring flood. Annual runoff coefficient is 0.32, and that for the flood period is ~0.56. Fossil flora 9 is derived from a loam and clay sediment layer in core SV-1–8, which is correlated with the Younger Dryas on the basis of pollen composition and radiocarbon dating. This horizon is distinguished by high nonarboreal pollen content, as well as a distinct maximum of Artemisia pollen and a smaller peak of Chenopodiaceae. Among trees, Pinus sylvestris and Betula alba were predominant. Of the dark coniferous trees, Picea abies and Pinus sibirica are registered in this flora. Of the cold-tolerant shrubs, Betula humilis and Alnaster are present. Relatively thermophilic trees once again disappear from the local flora (see Table 3). Changes in the flora are indicative of both colder and more arid climatic conditions compared to the previous time interval, with probable reestablishment of permafrost in the region. A region analogue for fossil flora 9 is situated in the middle reaches of the Biya River in the Altai Mountains (Fig. 6). In the vegetation cover of this area, open larch woodlands with steppe elements and patches of steppes are distributed next to fir and Siberian pine forests on the mountain slopes. Birch and Scots pine communities occur on sandy soil. The area is characterized by a cold, continental, and relatively dry climate with mean annual precipitation of 450 mm and runoff depth of 150–300 mm. The mean January air temperature is –18 °C, and that of July is 15 °C. Floras 10–12, deriving from palynological data of peat in core SV-1–8 from the Svapa River valley, indicate a conspicuous change of landscape and climatic conditions at the transition from the Late Glacial to the Holocene. These floras are dominated by forest and meadow plants (see Table 3). Their closest present-day analogues are located within the boreal forest zone and, for the Atlantic period of the Holocene, broad-leaved forest zone near the boundary of the steppe zone (see Fig. 6). These changes suggest a considerable rise in temperature and precipitation, decrease in the runoff coefficient because of degradation of permafrost, and, therefore, a substantial decrease in surface runoff (see Table 4). On the whole, changes in geographical position and hydroclimatic characteristics of the paleofloristic region analogues reflect a complexity of climate change during the Late Glacial against the background tendency toward warmer and wetter climate (see Fig. 7). Secondary oscillations can be seen on this general trend. Air temperature was the lowest in the LGM, in the late part of the Oldest Dryas, and in the Younger Dryas. Relatively warm intervals precede the Oldest Dryas and correspond to the Bølling and the Allerød. Extremely low winter temperatures

Surface runoff to the Black Sea from the East European Plain

19

indicate the existence of permafrost during the Late Pleniglacial, the Oldest Dryas to early Bølling, and in the Younger Dryas. Precipitation changes in the LGM and LGT generally followed temperature changes, the cold stages being relatively dry, and the warm stages relatively humid, with the maximum precipitation achieved in the pre–Oldest Dryas warming (ca. 14.5–14 14C k.y. B.P.), and in the Bølling-Allerød interstade (ca. 12.5–11 14C k.y. B.P.). Surface runoff is not determined entirely by annual precipitation, since water losses depend on air temperature. The runoff coefficient is inversely correlated with temperature changes, showing a strong overall decrease during the LGM-LGT. Generally following changes in runoff coefficient, the surface runoff also decreased considerably during this period (Fig. 7). Secondary oscillations of precipitation and water losses led to significant variations in surface runoff. The greatest runoff was characteristic of the pre–Oldest Dryas warming, because of relatively high precipitation, low losses, and preserved permafrost. The second highest runoff maximum was achieved at the beginning of the Bølling interval, when the temperatures still remained low while precipitation began to rise. During the cold and dry Younger Dryas, an increase in runoff was caused by low water losses to evaporation. During the warm and relatively humid Bølling and Allerød interstadials, higher losses to evaporation caused a relative decrease in surface runoff. Nevertheless, the entire LGM-LGT interval was characterized by high surface runoff compared to the present-day level, which was achieved at the beginning of the Holocene (see Fig. 7).

was assumed to be constant and equal to 60 mm yr–1, according to measurements on recent glaciers. The results are more qualitative than quantitative because of the great uncertainty about the structure and coefficients of Equations 1–8. Nevertheless, these estimates show a very large meltwater supply to the headwaters of the Dnieper River and its tributaries (Berezina and Pripyat Rivers), which at the early stages of deglaciation was comparable with recent annual surface runoff from the Dnieper River basin: 54 km3. The sizes of the blow-out channels coeval with this large meltwater supply (see Fig. 2) yield an additional possibility for estimating its volume. Using Equation 9 along with the width W and meander half-wavelength λ of these paleochannels (Table 5), the mean annual discharge Q and annual yield V were estimated. Presumably, the seasonality of runoff was high, so index of runoff variability y = 5. Paleochannel effective width W* was calculated as

Estimation of the Meltwater Supply in the Dnieper River Basin during the Early Stages of the Ice-Sheet Retreat

Estimation of Surface Runoff in the Dnieper and Don River Basins during the Period of Large Paleochannel Formation

Morphometrical parameters of the Fennoscandian ice sheet at the LGM (Bologoye)–Vepsovo stages (see Fig. 2) along with Equations 1–8 provide an estimate of change in glacier volume, snow accumulation, and final mass budget (Table 1). Such estimates were made for three main scenarios: a “cold” glacier with a shape similar to those of the Antarctic and Greenland glaciers; a “warm” glacier with a flatter hypothetical shape, derived from some assumptions about Fennoscandian glacier dynamics (Khodakov 1982); and a “dead-ice” glacier with a marginal zone of “inactive” ice with 500 m thickness, which follows from geomorphologic considerations (Chebotareva and Faustova, 1982; Faustova, 1984). Mean annual evaporation from the glacier surface

Calculations with Equations 9–11 require knowledge of paleochannel morphology and a modern hydrological analogue of the ancient river. Paleochannel effective width W* was estimated with Equation 12 with the mean relative error of 15% (see Table 2). The climatic region analogues for Late Glacial time were estimated from paleofloristic analysis (see Fig. 6). For relatively mild periods of the Late Glacial (pre–Oldest Dryas warming, Bølling and Allerød interstadials), analogues are situated in the basins of the Biya River (the Altai Mountains) and of the Belaya River (southeastern Ural Mountains). For the cold periods (the LGM, the Oldest, Older, and Younger Dryas), the modern region analogues are located in the southeastern sub-Baikal

W* =

W + (λ / kλ ) . 2

(12)

Here, kλ is λ/W; the mean value of the ratio is equal to 5.6 for the paleochannels in the plains of northern Eurasia (Sidorchuk et al., 2008). This estimate of annual meltwater supply (from 60 km3 annually) is close to the “dead-ice” scenario and in general confirms a very large volume of meltwater flowing to the Black Sea basin in the early stages of the ice-sheet retreat.

TABLE 5. MELTWATER FLOW THROUGH THE MAIN PALEOCHANNELS FROM THE ICE-DAM LAKES IN THE UPPER DNIEPER RIVER BASIN Meander Channel Effective Mean annual width, W width, W* discharge, Qa Paleochannel half-wavelength, λ 3 (m /s) (m) (m) (m) Paleochannel from the Upper Neman ice-dam lake to the Pripyat' River through the Schara River 5250 1600 1270 650 Paleo–Berezina River near Borisov 3540 1200 920 400 Paleo–Dnieper River near Rogachev 8600 1800 1670 950

Annual water yield volume, V 3 (km )

20 13 30

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region and at the headwaters of the Yenisei River (the Sayan Mountains). These territories can be used as hydrological analogues only partly (for example, for the estimation of runoff coefficient) because of the mountainous relief, high inclinations, and small areas of the river basins. For the estimation of such characteristics as discharge variability, lowland territories with similar climatic indexes (mean temperature of January and July, annual precipitation) and landscapes, close to those of the LGT (sparse open vegetation, widespread permafrost), must be selected. The lowland with climatic characteristics similar to those in the Altai and Ural Mountains, is the Bol’shezemel’skaya Tundra in the northeastern part of European Russia. The lowland with climatic characteristics similar to the Yenisei River headwaters is located within the Lena Plateau in central Yakutia (see Fig. 6). Hydrological regimes of large rivers in these two territories are quite similar: spring flood is high and sharp, and most of the annual flow passes during the flood. Low water period during the summer, autumn, and winter is long, but some years only 10%–12% of the annual runoff passes during this period. Coefficient a in Equation 11, which depends on hydrological regime, is therefore very similar for these two territories: it is 2.25 for the Bol’shezemel’skaya Tundra and 2.23 for the Lena Plateau. The sequence of paleohydrological calculations is as follows. The coefficient of within-year flow variability y for the past hydrological regime was estimated with Equation 11 for each river basin in Table 2, using coefficient a = 2.25. Basin area was assumed to be constant since the LGT. The mean annual paleodischarge Qa m3 s–1 was calculated with Equation 9, using y and effective width W*. The mean maximum paleodischarge Qmax

was estimated with Equation 10, using y and Qa. Surface runoff depth X (mm yr–1) from basin area F (km2) was calculated as

X = 31,536

Qa . F

(13)

The structure of Equation 9 leads to an increase in the relative error of discharge estimation compared to the relative error of channel width estimation. Therefore, the errors of paleodischarge and runoff depth estimates for the Dnieper and Don River basins are close to ±20%. Although we obtained only a few 14C dates from paleochannel deposits in the Dnieper and Don River basins (see descriptions of the key sites), these dates support a hypothesis of synchronous activity of the large rivers. As the oldest large paleochannels were abandoned 15–13 14C k.y. B.P., the large rivers were active in pre–Oldest Dryas time. Therefore, information from Table 2 can be used to reconstruct the spatial distribution of surface runoff within the Dnieper and Don River basins for this period (Fig. 8), and to estimate the volume of water flowing into the Black Sea from these basins. The general pattern of the pre–Oldest Dryas runoff is quite similar to the modern longitudinal decrease in runoff depth from north to south in the Dnieper and Don River basins (Figs. 8A and 8B). Maximum runoff was reconstructed for the area adjacent to the ice sheet, though none of the rivers used in the calculations was fed by glacier meltwater in the pre–Oldest Dryas. There was some lateral differentiation in the runoff: the Dnieper River basin was drier than the Don River basin. The same effect can be

Figure 8. Surface runoff in the East European part of the Black Sea drainage basin: (A) recent (measured) (after Evstigneev, 1990); and (B) calculated for the period of activity of large meandering rivers (pre–Oldest Dryas time). Key: (1) annual depth of surface runoff (mm); (2) data sites.

Surface runoff to the Black Sea from the East European Plain

River Don Dnieper Total

TABLE 6. VOLUME OF ANNUAL SURFACE RUNOFF FROM THE DNIEPER AND DON RIVER BASINS Recent Late Pleniglacial Annual runoff volume, VR Basin area Annual runoff volume, VP Basin area 2 3 2 3 (km ) (km ) (km ) (km ) 422,000 29 422,000 110 504,000 54 504,000 166 926,000 83 926,000 276

distinguished in the recent runoff pattern (Fig. 8A). The runoff during the pre–Oldest Dryas warming was much greater than the modern value. Runoff depth reached 600 mm in the basins of the Upper Dnieper and Don. It was also ~600 mm in the western part of the Dnieper River basin. In the eastern part of the Dnieper River basin, runoff depth decreased rapidly in the northsouth direction. It was ~400–600 mm at the headwaters and did not exceed 100–200 mm in the middle and lower Dnieper River basin. Runoff was more than 200 mm in the upper and middle Don River basin, and only at the lower reaches did it decrease to 100 mm. The excess of annual water flow above the modern value can be explained by greater precipitation and lower losses (greater runoff coefficient values). The rate of north-to-south decrease in runoff was lower in the past than at present. In the Dnieper River basin, the modern ratio between the runoff depth at the headwaters (200 mm) and lower reaches (20 mm) is 10; during the pre–Oldest Dryas warming, it was 6 (600 mm/100 mm). In the Don River basin, these ratios are closer to one another (150 mm/20 mm and 600 mm/100 mm) because of a smaller north-south extent of the basin. The map in Figure 8B allows one to calculate an annual volume of surface runoff for the period of large river activity and to compare it with recent characteristics (Table 6). During the pre– Oldest Dryas warming, the Don River supplied about four times the present-day water volume, and the Dnieper River supplied about three times the volume. If we apply these proportions to other rivers, the total annual surface runoff from the East European part of the Black Sea basin can be estimated at ~365 km3. Even if we keep recent values for the runoff from region III, the estimate will be more than 300 km3, i.e., almost three times greater than the modern runoff from this area and even greater than the modern runoff from the entire Black Sea drainage basin.

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VP/VR 3.8 3.1 3.3

that the meltwater supply was at least equal to recent runoff from the Dnieper River basin, and possibly in excess. This estimate is confirmed by the morphology of the blow-out channels, through which the meltwater passed. Paleoclimatic reconstructions for the LGM are based on data from the Drichaluki site, which characterize the Upper Dnieper basin. They reveal a cold and relatively dry environment with precipitation lower than at present (see Table 4). Low winter and summer air temperatures and the existence of deep permafrost led to low precipitation losses from infiltration and evaporation. Therefore, surface runoff was higher than recently, in spite of the lower precipitation. During the Late Pleniglacial, landscape and climatic conditions over the Russian Plain were less differentiated than at present, and the so-called periglacial hyperzone was formed (Velichko, 1973). At the LGM, permafrost in the Russian Plain reached 45°N (Velichko et al., 1982), thus including the major part of the Dnieper River basin. Taking this into consideration, annual meteoritic water runoff from the Dnieper basin on the whole was estimated at 100 km3, which is about two times greater than recent totals. With the additional 60 km3 of meltwater runoff, the cumulative runoff to the Black Sea from the Dnieper River basin at the LGM was significantly larger than recent runoff from the entire East European part of the Black Sea drainage basin. Extraordinary morphological features—large meandering paleochannels—mark the next hydrological period at the end

DISCUSSION Morphological, geological, geochronological, and palynological information on the East European part of the Black Sea drainage basin enables a reconstruction of the landscape, climate, and hydrological history of this region (Fig. 9). The amount of available paleogeographical data on different time intervals varies, influencing the reliability of the suggested interpretations. The weakest evidence is related to the hydrological regimen at the LGM. The great uncertainty in the paleoglaciological reconstruction of the shape and budget of the Fennoscandian ice sheet makes the estimate of meltwater supply to the Dnieper River basin more qualitative than quantitative. It is possible to say

Figure 9. Reconstructed surface runoff in the East European part of the Black Sea drainage basin during the Last Glacial Maximum (LGM) and Late Glacial time (LGT). PB—Preboreal; DR3—Younger Dryas; AL—Allerød; BØ—Bølling; DR1—Oldest Dryas.

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of the Pleniglacial (pre–Oldest Dryas warming). Existing dates (see Table 1 in Sidorchuk et al., 2008) show that this period had begun before 15 ka (the Moskva River paleochannel). In the Dnieper and Don River basins, widths and meander wavelengths of the large paleochannels are ten to fifteen times greater than those of the modern channels with the same drainage areas. The closest modern hydrological analogue for this period is an open tundra landscape with permafrost and long, severe winters with considerable snowfall (see Fig. 6). In such landscapes, short and high spring floods with large discharges formed large meandering channels. During the rest of the year, these channels were nearly empty. Using the large paleochannel morphology, we can roughly calculate the annual surface runoff from the East European part of the Black Sea basin during the pre–Oldest Dryas warming at ~300–365 km3. Because of low differentiation of landscapes within the periglacial hyperzone, we can assume that past/recent runoff ratios, estimated in the region analogues, can be extended to the whole East European part of the Black Sea basin. This assumption is supported by the results of our reconstructions. Thus, in the region analogue located for fossil flora 2 (see Fig. 6), the ratio between the past and recent runoffs is equal to 3.5 (see Table 4). A very similar ratio (3.2) was reconstructed for the Late Pleniglacial in the Moskva River basin near Moscow (Sidorchuk et al., 2009). Therefore, using the reconstructed values of the main hydroclimatic indexes for various intervals of the LGT (see Table 4), we calculated water discharge from the East European part of the Black Sea drainage (Fig. 9). Following the maximum reached at the end of the Pleniglacial, the discharge generally decreased, although it remained greater than the modern value until the beginning of the Holocene. The decrease in water discharge was interrupted by two events of high runoff. A secondary system of large paleochannels in the Khoper River basin (see Fig. 5) was formed during the second of these events, which corresponds to the Younger Dryas. The East European part of the Black Sea basin makes up about one third of its recent basin. Of the rest of the Black Sea basin, the Danube River basin is the largest part. When the Caspian Sea was connected to the Black Sea through the Manych Strait, the Black Sea drainage basin increased dramatically. Next, we shall discuss briefly the water supply from the Danube River and from the Caspian Sea during the LGM and LGT. The Danube River basin can be divided into two main morphological regions: (1) the Alpine and subalpine region of glacial morphology, which can be used for paleoglaciological reconstruction, and (2) the sub-Carpathian region, which shows clear traces of large river activity. Within the latter, the morphology of the large paleochannels can be used for paleohydrological reconstructions. The Last Glaciation in the eastern Alps, where most of the Danube tributaries have their origin, was described in detail by van Huzen (2004). Application of Equations 1–8 to van Huzen’s estimates of the position and age of various stages of the Last Glaciation gives the following results: during the period between ca. 21 and ca. 17 14C k.y. B.P., meltwater supply to the Upper Danube was very low; it was much higher (~35 km3 yr–1) during

the short period when glaciers retreated very rapidly between ca. 17 and ca. 16.5 14C k.y. B.P., and afterward, it gradually decreased until the beginning of the Holocene. The morphology and deposits of the large paleorivers were investigated by Borsy and Felegyhazi (1983), by Kasse et al. (2000) in Hungary, and by Howard et al. (2004) in Romania. These investigations showed that large rivers in the Tisza River basin were active at the end of the Pleniglacial, i.e., at the beginning of the LGT (Dr. K. Kasse, 2008, personal commun.). Large rivers in Romania were active before 12 14C k.y. B.P. (Howard et al., 2004). These estimates correspond in general to the dates obtained in the East European Plain (Sidorchuk et al., 2008). The widths and meander lengths of radiocarbon-dated large paleochannels were measured on Landsat-7 images. Calculations using Equations 9–11 show that, during the LGT, the Carpathian tributaries of the Danube supplied three times their present-day discharge. Presumably, the runoff from the Danube River basin changed in similar ways to that in the East European part of the Black Sea basin (see Fig. 9), with its maximum shifted toward the beginning of the LGT. The history of the last connection between the Black Sea and the Caspian Sea through the Manych Strait has been analyzed by many researchers (e.g., Goretskiy, 1957; Popov, 1983; Menabde and Svitoch, 1990), but still many questions remain unsolved, mostly regarding the chronology of events and the water budget. This connection occurred at some time during the early phase of the Khvalynian transgression of the Caspian Sea, which is dated to 20–7 14C k.y. B.P. (Svitoch, 1991). The maximum level of the first stage of this transgression was about +50 m asl or +78 m above the recent level of the Caspian Sea. This stage is marked by +50 m marine terraces with the deposits containing the Khvalynian malacofauna, and by the river terraces of the tributaries of the Lower Volga River. Alluvium that built up these terraces corresponds to marine and estuarine deposits of the transgression (Obedientova, 1977). The remnants of the large paleochannels on these river terraces (Sidorchuk et al., 2009) show that their formation took place within the stage of large river activity at the LGT, which corresponds to the latest radiocarbon dates associated with the maximum phase of the Khvalynian transgression: 14– 11 ka (Svitoch and Yanina, 1997), 16–11 ka (Leonov et al., 2002), 16–13 ka (Chepalyga et al., 2008). Our calculations for that stage of the Caspian Sea water budget (Panin et al., 2005; Sidorchuk et al., 2009) show that runoff from the large rivers in the Volga River basin (~500 km3 yr–1) was sufficient to build up the highest level of the Khvalynian transgression. However, this runoff was not large enough to maintain the long-term flow of Caspian water into the Black Sea through the Manych Strait. The mean sizes of fluvial features in the Manych Straight (channel width ~8 km, length of alternating bars ~20 km, height 15–20 m, and meander half-wavelength ~40 km) support the hypothesis of a very high discharge during their formation. The mean annual discharge of ~65,000 m3 s–1 calculated for a stream with such geometry using Equation 9 and y = 100 (stable runoff from a large lake, when Qmax is close to Qa) is greater than the flood discharge at the Lower Mississippi River. Annual volume of runoff through the Manych

Surface runoff to the Black Sea from the East European Plain Strait was ~2000 km3 yr–1. Note, that Chepalyga (2007) used a different method of discharge estimation: the product of channel cross-section area and flow velocity. For the Zunda-Tolga profile, cored by Popov (1983), the cross-section area was estimated as 250,000 m2 (width of 10,000 m on a mean depth of 25 m). Flow velocity was estimated at 0.2 m s–1 according to sediment particle size. Calculated discharge (Chepalyga, 2007) was 50,000 m3 s–1. That is rather close to our estimate, taking into account the completely different approach to figuring the result. The most probable hypothesis that can explain these facts is that a high threshold with its top at about +50 m asl existed in the middle of the Manych Strait. This natural dam separated the high-standing Caspian Sea from the relatively low-standing Black Sea (Menabde and Svitoch, 1990). At some moment, the dam was eroded to a level of about +22 m asl. As a result, the Caspian Sea water from between +50 and +22 m asl (~23,000 km3) was emptied into the Black Sea during a period of only 20–30 yr. This event happened within the large river stage. This assumption is supported by the existence of remnants of large paleochannels on the terraces of the tributaries of the Volga River, which were formed both before and after incision of those tributaries due to the Caspian Sea level drop (Sidorchuk et al., 2009). Some radiocarbon dates indicate that the flow via the Manych Strait took place between the beginning of the late Valdai/late Weichselian deglaciation and 12 ka (Arslanov and Yanina, 2008). Our reconstruction of high river runoff into the Black Sea at the LGM and LGT (2–2.5 times greater than the present; see Fig. 9) does not support the “flood hypothesis” proposed by Ryan et al. (1997). This hypothesis holds that the –120 to –150 m stage of the Black Sea lasted until ca. 7 14C k.y. B.P. The estimated high river runoff does not agree with a low level and isolated-lake status of the Black Sea throughout the LGM and LGT. Today, the Black Sea has a positive water balance with some 300 km3 of excess water removed annually via the Bosphorus (Özsoy et al., 1995). To maintain a low stage at the LGM and LGT, the values of effective evaporation should have been much higher than today. Reconstructions of Tarasov et al. (1999) for the LGM provide quite the opposite result, suggesting a rate of evaporation that was lower than present. Low evaporation aided by high water supply from the drainage basin make it more probable that the Black Sea was filled up to the Bosphorus sill (now ~35 m below sea level) in the LGT, with large amounts of excess freshwater flowing via this channel straight into the Marmara Sea. This is supported by recent findings of Giosan et al. (2009) that the subaerial deltaic plain of the Danube River, which is indicative of the sea-level position, was located around 30 m below sea level directly before the Black Sea to world ocean reconnection was established at ca. 8.4 ka. CONCLUSIONS Three time intervals can be distinguished within the general sequence of hydrological events in the East European part of the Black Sea drainage basin during the LGM and LGT.

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1. A cold and dry interval with precipitation lower than today and a high runoff coefficient caused by low evaporation losses and continuous spread of permafrost correlates with the LGM and the beginning of deglaciation, 18–15 14C k.y. B.P. A relatively high surface runoff (210 km3 yr–1) caused by precipitation was then supplemented by meltwater influx (60 km3 yr–1) from proglacial lakes. 2. A warmer, humid interval with precipitation exceeding that of today and a high runoff coefficient caused by relatively low temperatures correlates with the pre–Oldest Dryas time, 15–14 14C k.y. B.P. This was a period of very high surface runoff (300–365 km3 yr–1), when large meandering river channels were formed. A short event occurred within this period when a large amount of Caspian water (23,000 km3) was supplied to the Black Sea through the Manych Strait. 3. An interval with short-term climatic oscillations and nonsteady decrease of surface runoff, from ca. 14 to 10 14C k.y. B.P., corresponds to the LGT. Secondary climatic oscillations during the LGT were expressed against the overall trend toward warming, primarily the rise of winter temperatures. Phases with the lowest air temperatures (the Oldest and Younger Dryas) were favorable for permafrost development and to low losses, providing high runoff coefficients and, therefore, increasing the runoff (up to 355 km3 yr–1) despite a decrease in precipitation. Phases with relatively high air temperatures (Bølling and Allerød interstadials) were favorable for permafrost degradation and higher losses to evapotranspiration, providing low runoff coefficients and therefore decreasing the runoff (to 165 km3 yr–1) despite an increase in precipitation. By the end of this time interval, large meandering paleochannels were transformed into smaller channels. The entire LGM-LGT interval was characterized by high surface runoff from the East European part of the Black Sea drainage basin compared to that of the present day (110 km3 yr–1), which was achieved at the beginning of the Holocene. ACKNOWLEDGMENTS Financial support for this work was received from the Russian Foundation of Basic Research (RFBR), project no. 97-05-64708, 00-05-64021, and 09-05-00340. Valuable suggestions by reviewers J. Herget and T. Kalicki were incorporated into the text. REFERENCES CITED Aksu, A.E., Hiscott, R.N., Yaşar, D., Işler, F.I., and Marsh, S., 2002, Seismic stratigraphy of late Quaternary deposits from the southwestern Black Sea shelf: Evidence for non-catastrophic variations in sea-level during the last ~10,000 yr: Marine Geology, v. 190, p. 61–94, doi: 10.1016/S0025 -3227(02)00343-2. Arslanov, Kh., and Yanina, T., 2008, Radiocarbon age of the Khvalynian Manych passage, in Gilbert, A.S., and Yanko-Hombach, V., eds., Extended Abstracts of the Fourth Plenary Meeting and Field Trip of IGCP 521 “Black Sea–Mediterranean Corridor during the Last 30 k.y.: Sea Level Change and Human Adaptation,” and INQUA 0501 “Caspian–Black Sea– Mediterranean Corridor during the Last 30 k.y.: Sea Level Change and Human Adaptive Strategies” (4–16 October 2008; Bucharest, Romania, and Varna, Bulgaria): Bucharest, GeoEcoMar, p. 15–18.

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Svezhentsev, Y.S., 1993, Radiocarbon chronology for the Upper Paleolithic sites on the East European Plain, in Soffer, O., and Praslov, N.D., eds., From Kostenki to Clovis: New York, Plenum Press, p. 23–30. Svitoch, A.A., 1991, Kolebaniya urovnya Kaspiyskogo morya v pleistotsene [Fluctuations of the Caspian Sea level in the Pleistocene], in Scherbakov, F.A., and Svitoch, A.A., eds., Paleogeografiya i Geomorfologiya Kaspiyskogo Regiona v Pleistotsene [Palaeogeography and Geomorphology of the Caspian Region in the Pleistocene]: Moscow, Nauka, p. 5–100 (in Russian). Svitoch, A.A., and Yanina, T.A., 1997, Chetvertichnye otlozhenia poberezhiy Kaspiyskogo morya [Quaternary Sediments on the Caspian Sea Coasts]: Moscow, Rosselkhozacademia Press, 270 p. (in Russian). Szafer, W., 1946–1947, Flora plioceńska z Krościenka nad Dunajcem (The Pliocene flora of Kroscienko on Dunajec River): Rozprawy Wydzialu matemat.-przyrodn: Polska Akad. Umiejet., Krakow, 72 B, 1, 162 p.; 2, 213 p. (in Polish). Tarasov, P.E., Peyron, O., Guiot, J., Brewer, S.V., Volkova, V.S., Bezusko, L.G., Dorofeyuk, N.I., Kvavadze, E.V., Osipova, I.M., and Panova, N.K., 1999, Last Glacial Maximum climate of the former Soviet Union and Mongolia reconstructed from pollen and plant macrofossil data: Climate Dynamics, v. 15, p. 227–240, doi: 10.1007/s003820050278. van Huzen, D., 2004, Quaternary glaciations in Austria, in Ehlers, J., and Gibbard, P.L., eds., Quaternary Glaciations—Extent and Chronology: Amsterdam, Elsevier, p. 1–13. Varuschenko, S.I., Varuschenko, A.N., and Klige, R.K., 1987, Izmeneniya rezhima Kaspiyskogo morya i besstochnykh vodoyemov v paleovremeni [Changes of the Caspian Sea and Other Closed Lake Regimes in PalaeoTime]: Moscow, Nauka, 240 p. (in Russian). Velichkevich, F.Yu., 1982, Pleistotsenovye Flory Lednikovykh Oblastei Vostochno-Evropeiskoi Ravniny [Pleistocene Floras of the Glaciated Districts of the East European Plain]: Minsk, Nauka i Tekhnika, 239 p. (in Russian). Velichko, A.A., 1973, Prirodnyi protsess v pleistotsene [Natural Process in the Pleistocene]: Moscow, Nauka, 256 p. (in Russian). Velichko, A.A., Berdnikov, V.V., and Nechaev, V.P., 1982, Rekonstruktsiya zony mnogoletnei merzloty i etapov eyo razvitiya [Reconstruction of the permafrost zone and stages of its development], in Gerasimov, I.P., and Velichko, A.A., eds., Paleogeografiya Evropy za poslednie sto tysyach let (Atlas-monografiya) [Palaeogeography of Europe during the Last 100,000 Years (Atlas-Monograph)]: Moscow, Nauka, p. 74–81 (in Russian). Velichko, A.A., Faustova, M.A., Gribchenko, Yu.N., Pisareva, V.V., and Sudakova, N.G., 2004, Glaciations of the East European Plain—Distribution and chronology, in Ehlers, J., and Gibbard, P.L., eds., Quaternary Glaciations—Extent and Chronology: Amsterdam, Elsevier, p. 337–354. Vlasov, V.K., and Kulikov, O.A., 1988, Radiotermolyuminestsentnyi metod datirovaniya rykhlykh otlozheniy [Radio-thermo-luminescence method of Non-Consolidated Sediment Dating]: Moscow, Moscow State University Press, 72 p. (in Russian). Volkov, I.A., 1960, O nedavnem proshlom rek Ishim i Nura [On the recent past of the rivers Ishim and Nura]: Trudy Laboratorii Aerometodov AN SSSR [Proceedings of the Laboratory of Aeromethods, USSR Academy of Sciences], v. 9, p. 15–19 (in Russian). Volkov, I.A., 1963, Sledy moschnogo stoka v dolinakh yuga Zapadnoi Sibiri [An evidence of the powerful flow in the valleys of west Siberian rivers]: Doklady AN SSSR [Reports of the USSR Academy of Sciences], v. 151, p. 648–651 (in Russian). Walker, M.J.C., Björck, S., Lowe, J.J., Cwynar, L.C., Knudsen, K.-L., Wohlfarth, B., and INTIMATE Group I, 1999, Isotopic ‘events’ in the GRIP ice core: A stratotype for the late Pleistocene: Quaternary Science Reviews, v. 18, p. 1143–1150. Winguth, C., Wong, H.K., Panin, N., Dinu, C., Georgescu, P., Ungureanu, G., Krugliakov, V.V., and Podshuveit, V., 2000, Upper Quaternary water level history and sedimentation in the northwestern Black Sea: Marine Geology, v. 167, p. 127–146, doi: 10.1016/S0025-3227(00)00024-4. Zagwijn, W.H., 1996, An analysis of Eemian climate in Western and Central Europe: Quaternary Science Reviews, v. 15, p. 451–469, doi: 10.1016/0277-3791(96)00011-X.

MANUSCRIPT ACCEPTED BY THE SOCIETY 22 JUNE 2010 Printed in the USA

The Geological Society of America Special Paper 473 2011

Modeling extreme Black Sea and Caspian Sea levels of the past 21,000 years with general circulation models Alexander Kislov Pavel Toropov Department of Meteorology and Climatology, Faculty of Geography, M.V. Lomonosov Moscow State University, Leninskiye Gory, Moscow 119992, Russia

ABSTRACT This paper describes the relationship between sea levels and climate based on the links between sea-level variations and river runoff. During the final late Pleistocene and postglacial periods, the Caspian Sea fluctuated between regression and transgression stages. The Black Sea experienced fluctuations as well, but these were mainly controlled by the world ocean due to water exchange through the Bosporus Strait. Sometimes, the Caspian Sea overflowed into the Black Sea through the Manych Strait, and they periodically coalesced. Change in the level of both seas could be interpreted as responses to the regional-scale water budget (the balance between inflow and outflow components). These components can be calculated from atmospheric general circulation models. This approach uses climate modeling data to reproduce river runoff changes, and, consequently, variations in seawater and sea level under contrasting climate conditions. In response to glacial conditions of the last cold Pleistocene event, the lowering levels of the Black Sea (post-Karangatian regression stage) and the Caspian Sea (Atelian regression stage) are simulated simultaneously. This lends credence to the idea of the connection between deep regression states of the Caspian and Black Seas and mature stages of the late Quaternary glacial/cooling/drying planetary events. Analysis of observed information allows us to conclude—taking into account the uncertainties of reconstructed data—that at least two regression stages occurred simultaneously with late Quaternary glacial planetary events. The simulation of transgression stages (their onset and duration) remains a very difficult problem. Results of modeling have shown that during the warm periods (taking as examples the mid-Holocene and Allerød events), simulated river runoff did not increase to the extent needed for a strong transgression and overflow of the Caspian Sea into the Black Sea through the Manych Strait. Thus, there is no clear understanding about the source of “additional” water volume necessary to elevate the level of the Caspian Sea to a point that would permit overflow into the Black Sea.

Kislov, A., and Toropov, P., 2011, Modeling extreme Black Sea and Caspian Sea levels of the past 21,000 years with general circulation models, in Buynevich, I.V., Yanko-Hombach, V., Gilbert, A.S., and Martin, R.E., eds., Geology and Geoarchaeology of the Black Sea Region: Beyond the Flood Hypothesis: Geological Society of America Special Paper 473, p. 27–32, doi: 10.1130/2011.2473(02). For permission to copy, contact [email protected]. © 2011 The Geological Society of America. All rights reserved.

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INTRODUCTION The targets of our exploration are the Caspian Sea and the Black Sea (including the Sea of Azov). This paper describes the relationship between sea level and climate based on links between variations in sea level and the runoff from rivers. During the final late Pleistocene and postglacial periods, the Caspian Sea fluctuated between regressive and transgressive stages. The Black Sea experienced fluctuations as well, but these were mainly controlled by the world ocean due to water exchange through the Bosporus Strait. Sometimes, the Caspian Sea overflowed into the Black Sea through the Manych Strait (Chepalyga, 2007), and they periodically coalesced. The reasons for these sea-level changes could be different, and such a problem can be considered from the perspective of responses to the regional-scale water budget based upon planetary climate changes. The variability of seawater mass is a function of the balance between inflow and outflow components, and these factors are functions of the climate regime. Therefore, they could be calculated from atmospheric general circulation models (GCMs). This approach would allow the use of climate modeling data to reproduce river runoff changes, and, consequently, variations in seawater and sea levels under contrasting climate conditions. This paper is organized as follows: first, previous studies are reviewed; second, global climate changes during the Quaternary and their causes are presented; third, a description of the models is presented. The discussion then considers the closed-basin model that addresses the Caspian and Black Seas. The next sections present results from a seawater balance model that addresses the effects of drainage basin runoff changes and sea levels. The last section examines the paradox of “additional water.”

GLOBAL CLIMATE CHANGES DURING THE QUATERNARY The Quaternary has been characterized by both cold and warm phases. One candidate for a forcing agent that could produce such pronounced global climate variations is the Milankovitch mechanism (Berger, 1988). According to this theory, Earth’s orbital parameters change due to the influence of the Moon, Sun, and planets. Over 100,000 and 400,000 yr periods, eccentricity slowly varies, inducing small changes in the mean annual total insolation received by Earth. Obliquity oscillates from ~22° to ~25° over a 41,000 yr period, and the position of the equinoxes precesses relative to the perihelion with 19,000 and 23,000 yr periodicities. Obliquity and precession do not lead to global changes in mean annual energy but strongly modulate the seasonal pattern of insolation. The evidence for Milankovitch forcing of climate changes may be questioned for several reasons. First, there is no evidence that the climate cycles are periodic rather than aperiodic. Core records, both ice and deep sea, suggest that the dominant character is that of a random red-noise process. Second, much of the energy in low-frequency climate change occurs at periods around 100 ka, where the insolation forcing is very weak. The contribution of the Milankovitch periodicities (41,000, 19,000, and 23,000 yr periods) to climate change provides only a small fraction (15%–20%) of total climate variance. At times when orbital agents work synchronously, however, climatic response to variation in solar insolation can be distinguished from such noise. This effect led to the transition from the cold late Pleistocene to the warm Holocene, a transition that was not gradual but instead complicated by short-term events—e.g., the Allerød (AL)–Younger Dryas (YD) cycle. Many authors link the origin of these cycles to the behavior of the Atlantic thermohaline circulation.

PREVIOUS STUDIES RESULTS Various conceptual models have been used to link climate change to lake level (h), lake surface area (f), and catchment area (F) (e.g., Kalinin, 1968; Street-Perrot and Harrison, 1985; Benson and Paillet, 1989). Kalinin (1968) demonstrated that h asymptotically comes to an equilibrium level that is determined by the steady-state water-budget condition. Street-Perrot and Harrison (1985) classified closed lakes using functions relating f and F based on precipitation onto and evaporation from a lake surface and catchment (runoff) processes. Benson and Paillet (1989) argued that topographic constraints in lakes with more than one subbasin mean that lake area is the most appropriate measure of lake response to hydrologic balance. Few studies have been able to quantify the lake-climate relationship with precision for individual lakes. Functions of h, f, or precipitation and evaporation relationships have been used to evaluate climate model simulations (e.g., Kislov and Sourkova, 1998; Qin et al., 1998) and diagnose past climate (e.g., Harrison et al., 1993, 1996; Jones et al., 2001).

We studied the history of sea-level variation as influenced by climate conditions based on the results of climate modeling. Numerical experiments were used to understand how climatic parameters important to the Black and Caspian Sea water budget might change due to external forces. A modeling initiative, the Paleoclimate Modeling Intercomparison Project (PMIP), has focused on two slices of the past: (1) the mid-Holocene (6 ka calendar yr B.P. or ca. 5.3 ka radiocarbon yr B.P.), and (2) the last cold event of the late Quaternary (21 ka calendar yr B.P. or ca. 18 ka radiocarbon yr B.P.) because climatic conditions were remarkably different at those times, and abundant data are available that describe their environmental properties. In addition, both the Allerød (ca. 14.5 ka calendar yr B.P.) and the Younger Dryas (ca. 12 ka calendar yr B.P.) were studied as examples of short-scale variability. These slices (together with the modern state) are used in this paper to assess the linkage between climate variability and hydrological regime.

Modeling extreme Black Sea and Caspian Sea levels of the past 21,000 years A GCM (general circulation model) consists of an atmospheric model interactively coupled to submodels of the ocean, sea ice, land surface, and soil. Models have a typical horizontal resolution ~2°–4° latitude × longitude. Designations of the PMIP GCMs are presented in Kislov and Toropov (2007). Table 1 specifies all model boundary conditions and parameters. Closed-Basin Water-Budget Model and Data Quality We calculated changes in sea-surface area for each climatic event assuming that the closed sea was in hydrologic equilibrium with climate conditions. This is a reasonable assumption when one considers the impact of gradual climatic change on a sea with a short hydrologic response time compared to the typical rate of change for external forces. The steady-state equation for the annual water budget for a closed basin has the form: ef = YF,

(1)

where e = E – P, P is on-sea precipitation (m/yr), E is the evaporation (m/yr) per unit of lake area, Y is the runoff (m/yr) per unit of basin area, F is the drainage basin area, and f is the sea-surface area. Equation 1 assumes that the net groundwater flux into or out of the sea was probably minimal. Variation in lake area relative to the present status (denoted by index “0”) may be expressed in the form:

Δf ΔY ΔF Δe . = + − f0 Y0 F0 e0

(2)

This allows for the evaluation of the contribution of different factors toward the change in level (h) using information about lake size, bathymetry, and the surrounding topography as: Δh = (Δh)Y + (Δh)F + (Δh)e.

(3)

Evaluation of these factors can be undertaken using different approaches. Information about the change in catchment area

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(Δh)F can be extracted from paleogeographical data, or it can be calculated based on simulation of Earth’s surface paleotopography. Level change (Δh)Y + (Δh)e due to both runoff change and change in e is calculated based on data acquired through numerical simulations of GCMs. Under current climate conditions, the volume of water annually delivered by the rivers of the East European Plain to the Black Sea is 312 km3 (total runoff to the Black Sea is 338 km3). The largest rivers are the Danube, Dneiper, Southern Bug, and Don. Their contribution is ~90% of the total mean volume of river runoff. The river flow is determined by the balance between precipitation and evaporation; hence, river runoff responds immediately to climate changes. It is important that both precipitation and evaporation be calculated using a GCM. Moreover, on a large plain such as the East European Plain, GCM data reflect the state of the climate better than areas with a mosaic of complex surface conditions. The contribution of runoff represents a large fraction (~50%) of inflow volume. Contributions from sea-surface precipitation and inflow from the Sea of Marmara represent 30% and 20%, respectively. Sea evaporation is 430 km3 (~53% of outflow volume), and release of water to the Sea of Marmara comprises ~47% of the outflow volume. The Caspian Sea (a vast inland lake) is fed by several rivers. Their annual runoff from the East European Plain is 274 km3. The greatest contribution (more than 80% of the mean total volume of runoff) derives from the Volga River. Other principal components of the annual water budget are precipitation over the sea (which adds 76 km3) and evaporation from the sea surface (which removes 362 km3). A special concern for climate modeling is the quality of the information. The data presented in Kislov and Toropov (2007) indicate how well PMIP GCMs simulate today’s river runoff within the basins of the Caspian Sea and Black Sea. Results were considered “successful” only if the errors of modeled runoff volume for each basin lay within 20%, because this variability does not fall outside the limits of natural variability. According to this classification, the most “successful” GCMs were chosen. Data from these models were examined more closely during modeling of other climatic regimes.

TABLE 1. PALEOCLIMATE MODELING INTERCOMPARISON PROJECT BOUNDARY CONDITIONS AND PARAMETERS Time Time Boundary conditions and parameters Control experiment (6 ka B.P.) (21 ka B.P.) Sea-surface temperature and sea ice Modern Modern Calculated or prescribed by CLIMAP Continental ice sheets Modern M odern Prescribed (Peltier, 1994) Vegetation and land-surface characteristics Modern Modern Modern (besides areas covered by ice) Aerosol optical depth Modern M odern M odern –2 –2 –2 1365 Wm 1365 Wm Solar constant 1365 Wm Ecc = 0.018994 Ecc = 0.016724 Ecc = 0.018682 Orbital parameters ε = 23.446° ε = 24.105° ε = 22.949° λ = 102.04° λ = 0.87° λ = 114.42° 28 0 p p m 280 p pm 2 00 p pm CO2 Note: CLIMAP—Climate: Long-range Investigation, Mapping, and Prediction. Ecc—eccentricity; ε—obliquity; λ—longitude of the perihelion.

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Simulation of River Runoff and Sea-Level Changes during the Mid-Holocene Warm Event The Holocene has been marked by a relatively stable, warm climate that experiences weak global changes on average. By the early Holocene (9 ka), CO2 was up to its preindustrial level (Raynaud et al., 1993), the Scandinavian ice sheet was almost gone, and the Laurentide ice sheet had shrunk considerably (Dyke and Prest, 1987; Svensson, 1991). Sea-surface temperature (SST) was not significantly different from today. In the mid-Holocene, at 6 ka, insolation anomalies (compared to modern ones) were +5% in summer and −5% in winter, less than those at the beginning of the Holocene. PMIP simulations of the 6 ka climate were thus designed as pure sensitivity experiments to gauge changes of insolation forcing (see Table 1). Global model results were evaluated (Joussaume, 1999; Kohfeld and Harrison, 2000) against terrestrial proxy data (lake levels, pollen- and plant macrofossil–based reconstructions of 6 ka vegetation). Our calculations indicate that at 6 ka, there was no significant change in the East European Plain river runoff (see Table 2). The Volga River’s contribution to Caspian Sea inflow was slightly increased (93%) compared to today (88% based on PMIP model simulation and 84% based on observation). The value of (Δh)e over the Caspian Sea was estimated—based on simplified regional climate modeling (Kislov and Sourkova, 1998)—as a small value relative to the first term in Equation 2. Therefore, there was no large change in sea-surface area or level (this assessment is not true for the Black Sea because, at this time, it was not closed, and changes in its level were controlled by water exchange with the Mediterranean Sea). Simulation of River Runoff and Sea-Level Changes during the Last Cold Pleistocene Event The last cold event of the Pleistocene involved substantial changes in surface boundary conditions, and within the PMIP, a specific set was prepared (see Table 1). At 21 ka, the latitudinal distribution of insolation and its relative seasonal strength were similar to those of today (for example, in the northern summer, the solar energy deficit received by Earth was −2–4 Wm2). The ice-sheet extent and height were provided by Peltier (1994). CO2 concentration was estimated at 200 ppm as inferred from Antarctic ice cores (Raynaud et al., 1993). Over the oceans, two sets of PMIP experiments were defined: (1) prescribing

TABLE 2. CHANGE IN RIVER RUNOFF (Y – Y0)/Y0 (%) BELONGING TO THE BLACK AND CASPIAN SEAS AT 6 ka B.P. BASED ON PMIP DATA Black Sea Caspian Sea Ensemble of PMIP models +9 +14 –5 +5 Successful PMIP models Note: PMIP—Paleoclimate Modeling Intercomparison Project.

SST changes from estimates given by CLIMAP (1981), and (2) SST computed using coupled atmosphere–mixed-layer ocean models. Other kinds of intrinsic variability (e.g., due to ocean circulation changes) or other kinds of natural forces (e.g., solar irradiance and volcanic forces) were not incorporated into these experiments. Note that the boundary conditions used by the PMIP contain some uncertainties. For example, there is an underestimation of CLIMAP SST anomalies in the tropics (Anderson and Webb, 1994; Guilderson et al., 1994; Hostetler and Mix, 1999) and a problem with location and seasonal behavior of sea-ice cover (Weinelt et al., 1996). Another example touches upon the location of Quaternary ice sheets. Model sensitivity to these scenarios has been investigated previously (Kislov et al., 2002). It was shown that most differences (2–6 °C) between the results occurred within regions where the position of the ice sheet has changed. Thus, the difference in boundary condition provides only a regional effect. Moreover, taking into account the intermodel deviations in simulation data, these uncertainties are not crucial for global climate modeling. We evaluated the global model results against terrestrial proxy data (Joussaume, 1999; Kohfeld and Harrison, 2000; Kislov et al., 2002). These results clearly demonstrate that temperatures in the PMIP experiment reproduce the main peculiarities of reconstructed land temperature fields, but over the tropics, the simulations with prescribed CLIMAP SSTs produce too weak a cooling effect over land. All models produce drying in the extratropical zone, although the extent and location of the regions experiencing increased aridity vary between models. Consider the results of calculation of annual river runoff volumes for the Caspian Sea and the Black Sea (Table 3). At 21 ka, the total river runoff to the Caspian Sea (calculated by “successful” models) was substantially lower (–50%) compared to today. The relative contribution of Volga River runoff has a value of 72%. These facts are in accordance with the observational data. Taking into account that the second term in Equation 3 is equal to zero, and the value Δe/e0 can be estimated as small relative to the first term in Equation 2, we can estimate the relative decrease in Caspian Sea area as 50%, which means a substantial drop in level (~50 m). Calculated river runoff into the Black Sea decreased substantially as well (–45%), and the drop in level has been estimated at –200 m. During this period, there was no water exchange through the modern Bosporus Strait due to dropping of both the Black Sea level and the world ocean level.

TABLE 3. CHANGE IN RIVER RUNOFF (Y – Y0)/Y0 (%) INTO THE BLACK AND CASPIAN SEAS AT 21 ka B.P. BASED ON PMIP DATA Black Sea Caspian Sea Ensemble of PMIP models –22 –40 Successful PMIP models –45 –56 Note: PMIP—Paleoclimate Modeling Intercomparison Project.

Modeling extreme Black Sea and Caspian Sea levels of the past 21,000 years Simulation of River Runoff and Sea-Level Changes during the Cold Younger Dryas and the Warm Allerød Events At 12–14 ka, insolation anomalies were +5% in summer and −5% in winter (compared to mid-Holocene levels). During the YD, SST over the North Atlantic Ocean was significantly less than today, but during the Allerød, there were no large differences. For the climate simulation, GCM T42L15 was used with prescribed SSTs. Therefore, experimentally, the Allerød was practically similar to that discussed for the mid-Holocene experiment but with more active radiation forcing. The YD experiment yielded both strong radiation forcing and an SST anomaly in the North Atlantic Ocean. During the Allerød, annual river runoff volume into the Caspian and Black Seas was slightly increased (6%), and during the YD, it was lower (–12%) compared to today. Therefore, it could not cause serious sea-level changes. DISCUSSION The results of climate simulation presented here are important in light of the problem of chronologically correlating paleogeographical events that belong to different regional scales. In response to glacial conditions of the last Pleistocene cold event, the declining levels of the Black Sea (post-Karangatian regression stage) and the Caspian Sea (Atelian regression stage) are simulated simultaneously. Hence, these changes in sea level reflect climate forces at the planetary scale. This lends credence to the idea of a connection between the deep regression states of the Caspian and Black Seas and the mature stages of the late Quaternary glacial/cooling/drying planetary events. The next question is whether all deep regression stages exhibited by these seas have had a similar origin. Is the conclusion assigned to one snapshot at 21 ka similarly applicable to other events? This idea is probably true when sea-level curves denoting time-behavior of the Black and Caspian Seas are compared to the curve depicting global climate changes (Fig. 1). Taking into account uncertainties in the reconstructed data, it is possible to conclude that at least two final regression stages occurred simultaneously with late Quaternary glacial planetary events. As far as the transgression stages are concerned, the simulation of their onset and duration remains a very difficult problem. The aforementioned modeling results have shown that during the warm periods (taking as an example the mid-Holocene and Allerød events), the simulated river runoff did not increase sufficiently to create a strong Caspian Sea transgression leading to overflow into the Black Sea through the Manych Strait. In short, there is no evidence to connect the large, well-documented Khvalynian transgression stage of the Caspian Sea with the warm Allerød event. Thus, there is no clear understanding about the source of “additional” water volume capable of establishing a high enough Caspian Sea level that would permit overflow from the Caspian into the Black Sea. There are several speculative hypotheses that

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lie beyond the paradigm of climate modeling that attempt to solve this paradox. One idea is that Siberian rivers such as the Ob’, Yenisei, Lena, and others were bordered by ice sheets along the Arctic coast, and water from huge dammed lakes overflowed into the Caspian Sea through different spillways (Grosswald, 1998). A second hypothesis posits that the source of additional water may have been connected to an increase in precipitation over the Tian Shan Mountains due to penetration of the Indian monsoon during warm periods (Kislov and Toropov, 2007). Subsequently, meltwater outflow was directed from the Amu Daria toward the Uzboy Valley and the Caspian Sea throughout the intermediate Sarykamish Lake. However, there is no reliable evidence to support these hypotheses. The volume of river runoff would be effectively increased if the runoff coefficient (i.e., the ratio of runoff to precipitation) were to change. This coefficient increases, for example, if water from precipitation is not absorbed into the soil due to the presence of permafrost. There is some evidence that permafrost conditions existed within the East European Plain during the postglacial, and it is thought that the effect of such a scenario would have been to produce changes as large as 30% in runoff volume. ACKNOWLEDGMENTS Financial support for this work was provided by the Russian Fund for Basic Research. This research is a contribution to the PMIP (Paleoclimate Modeling Intercomparison Project) and IGCP (International Geological Correlation Programme).

Figure 1. Global climate change (marine isotope data, after Imbrie et al., 1984) and Black Sea and Caspian Sea level variations (after Shuisky [2007] and Svitoch [2003]).

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Kislov and Toropov

REFERENCES CITED Anderson, D.M., and Webb, R.S., 1994, Ice-age tropics revisited: Nature, v. 367, p. 23–24, doi: 10.1038/367023a0. Benson, L.V., and Paillet, F.L., 1989, The use of total lake-surface area as an indicator of climate change: Examples from the Lahontan Basin: Quaternary Research, v. 32, p. 262–275, doi: 10.1016/0033-5894(89)90093-8. Berger, A., 1988, Milankovitch theory and climate: Reviews of Geophysics, v. 26, p. 624–657, doi: 10.1029/RG026i004p00624. Chepalyga, A., 2007, The Late Glacial Great Flood in the Ponto-Caspian basin, in Yanko-Hombach, V., Gilbert, A.S., Panin, N., and Dolukhanov, P.M., eds., The Black Sea Flood Question: Changes in Coastline, Climate, and Human Settlement: Dordrecht, The Netherlands, Springer, p. 119–148. CLIMAP, 1981, Seasonal Reconstructions of the Earth’s Surface at the Last Glacial Maximum: Geological Society of America Map Series MC-36. Dyke, A.S., and Prest, V.K., 1987, Late Wisconsinan and Holocene history of the Laurentide ice sheet: Géographie Physique et Quaternaire, v. 41, p. 237–263. Grosswald, M.G., 1998, New approach to the Ice Age paleohydrology of northern Eurasia, in Benito, G., Baker, V.R., and Gregory, K.J., eds., Paleohydrology and Environmental Change: Chichester and New York, John Wiley & Sons, p. 199–214. Guilderson, T.P., Fairbanks, R.G., and Rubenstone, J.L., 1994, Tropical temperature variations since 20,000 years ago: Modulating interhemispheric climate change: Science, v. 263, p. 663–665, doi: 10.1126/ science.263.5147.663. Harrison, S.P., Prentice, I.C., and Guiot, J., 1993, Climatic controls on Holocene lake-level changes in Europe: Climate Dynamics, v. 8, p. 189–200, doi: 10.1007/BF00207965. Harrison, S.P., Yu, G., and Tarasov, P.E., 1996, Late Quaternary lake-level record from northern Eurasia: Quaternary Research, v. 45, p. 138–159, doi: 10.1006/qres.1996.0016. Hostetler, S.W., and Mix, A.C., 1999, Reassessment of ice-age cooling of the tropical ocean and atmosphere: Nature, v. 399, p. 673–676, doi: 10.1038/21401. Imbrie, J., Hays, J.D., Martinson, D.G., McIntyre, A., Mix, A.C., Morley, J.J., Pisias, N.G., Prell, W.L., and Shackleton, N.J., 1984, The orbital theory of Pleistocene climate: Support from a revised chronology of the marine δ18O record, in Berger, A., Imbrie, J., Hays, H., Kukla, G., and Saltzman, B., eds., Milankovitch and Climate: Understanding the Response to Astronomical Forcing, Proceedings of the NATO Advanced Research Workshop (held 30 November–4 December 1982, in Palisades, New York): Dordrecht, Netherlands, D. Reidel Publishing, p. 269–305. Jones, R.N., McMahon, T.A., and Bowlers, J.M., 2001, Modelling historical lake levels and recent climate change at three closed lakes, Western Victoria, Australia (c. 1840–1990): Journal of Hydrology (Amsterdam), v. 246, p. 159–180, doi: 10.1016/S0022-1694(01)00369-9. Joussaume, S., 1999, Modeling extreme climates of the past 20,000 years with general circulation models, in Holland, W.R., Joussaume, S., and David,

F., eds., Modeling the Earth’s Climate and its Variability: Amsterdam, Elsevier, p. 527–565. Kalinin, G.P., 1968, Problemy Globalnoi Gidrologii (Problems of Global Hydrology): Leningrad, Gidrometeoizdat, 376 p. (in Russian). Kislov, A.V., and Sourkova, G.V., 1998, Simulation of the Caspian Sea level changes during last 20000 years, in Benito, G., Baker, V.R., and Gregory, K.J., eds., Palaeohydrology and Environmental Change: Chichester, John Wiley & Sons, p. 235–246. Kislov, A.V., and Toropov, P.A., 2007, Climate modeling results for the circumPontic region from the late Pleistocene to the mid-Holocene, in YankoHombach, V., Gilbert, A.S., Panin, N., and Dolukhanov, P.M., eds., The Black Sea Flood Question: Changes in Coastline, Climate, and Human Settlement: Dordrecht, The Netherlands, Springer, p. 47–62. Kislov, A.V., Tarasov, P.E., and Sourkova, G.V., 2002, Pollen and other proxybased reconstructions and PMIP simulations of the Last Glacial Maximum mean annual temperature: An attempt to harmonize the data-model comparison procedure: Acta Palaeontologica Sinica, v. 41, p. 539–545. Kohfeld, K.E., and Harrison, S.P., 2000, How well can we simulate past climates? Evaluating the models using global palaeoenvironmental datasets: Quaternary Science Reviews, v. 19, p. 321–346, doi: 10.1016/S0277 -3791(99)00068-2. Peltier, W.R., 1994, Ice age paleotopography: Science, v. 265, p. 195–201, doi: 10.1126/science.265.5169.195. Qin, B., Harrison, S.P., and Kutzbach, J.E., 1998, Evaluation of modelled regional water balance using lake status data: A comparison of 6 ka simulations with the NCAR CCM: Quaternary Science Reviews, v. 17, p. 535–548, doi: 10.1016/S0277-3791(98)00011-0. Raynaud, D., Jouzel, J., Barnola, J.-M., Chappellaz, J., Delmas, R.J., and Lorius, C., 1993, The ice record of greenhouse gases: Science, v. 259, p. 926–934. Shuisky, Y., 2007, Climate dynamics, sea-level change, and shoreline migration in the Ukrainian sector of the circum-Pontic region, in Yanko-Hombach, V., Gilbert, A.S., Panin, N., and Dolukhanov, P.M., eds., The Black Sea Flood Question: Changes in Coastline, Climate, and Human Settlement: Dordrecht, The Netherlands, Springer, p. 251–277. Street-Perrot, F.A., and Harrison, S.P., 1985, Lake levels and climate reconstruction, in Hecht, A.D., ed., Paleoclimatic Analysis and Modeling: New York, John Wiley and Sons, p. 291–340. Svensson, N.O., 1991, Postglacial land uplift patterns of south Sweden and the Baltic Sea region: Terra Nova, v. 3, p. 369–378, doi: 10.1111/j.1365 -3121.1991.tb00165.x. Svitoch, A., 2003, Morskoi Pleistotsen Poberezhii Rossii (Marine Pleistocene of the Russian Coasts): Moscow, GEOS, 362 p. Weinelt, M., Sarnthein, M., Pflaumann, U., Schulz, H., Jung, S.J.A., and Erlenkeuser, H., 1996, Ice-free Nordic seas during the Last Glacial Maximum? Potential sites of deepwater formation: Paleoclimates, v. 1, p. 283–309.

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The Geological Society of America Special Paper 473 2011

Assessment of Black Sea water-level fluctuations since the Last Glacial Maximum G. Lericolais Institut Français de Recherche pour l’Exploitation de la Mer (IFREMER), Centre de BREST, BP 70, F29200 Plouzané cedex, France F. Guichard Laboratoire des Sciences du Climat et de l’Environnement (LSCE), CNRS-CEA, Avenue de la Terrasse, BP 1, 91198 Gif-sur-Yvette cedex, France C. Morigi Geological Survey of Denmark and Greenland (GEUS), Department of Stratigraphy, Øster Voldgade 10, 1350 Copenhagen, Denmark I. Popescu Institutul National de Cercetare-Dezvoltare pentru Geologie si Geoecologie Marina (GeoEcoMar), 23-25 Dimitrie Onciul Str, BP 34-51, Bucuresti, Romania C. Bulois School of Geological Sciences, University College Dublin, Belfield, Dublin 4, Ireland H. Gillet Unité Mixte de Recherche (UMR) 5805, Environnements et Paléoenvironnements Océaniques (EPOC), Université Bordeaux 1, Avenue des Facultés, F33405 Talence, France W.B.F. Ryan Lamont-Doherty Earth Observatory, Columbia University, 61 Route 9w, Palisades, New York 10964, USA

ABSTRACT This paper presents geophysical and core data obtained from several marine geology surveys carried out in the western Black Sea. These data provide a solid record of water-level fluctuation during the Last Glacial Maximum in the Black Sea. A Last Glacial Maximum lowstand wedge evidenced at the shelf edge in Romania, Bulgaria, and Turkey represents the starting point of this record. Then, a first transgressive system is identified as the Danube prodelta built under ~40 m of water depth. The related rise in water level is interpreted to have been caused by an increase in water provided to the Black Sea by the melting of the ice after 18,000 yr B.P., drained by the largest European rivers (Danube, Dnieper, Dniester). Subsequently, the Black Sea Lericolais, G., Guichard, F., Morigi, C., Popescu, I., Bulois, C., Gillet, H., and Ryan, W.B.F., 2011, Assessment of Black Sea water-level fluctuations since the Last Glacial Maximum, in Buynevich, I.V., Yanko-Hombach, V., Gilbert, A.S., and Martin, R.E., eds., Geology and Geoarchaeology of the Black Sea Region: Beyond the Flood Hypothesis: Geological Society of America Special Paper 473, p. 33–50, doi: 10.1130/2011.2473(03). For permission to copy, contact editing@geosociety .org. © 2011 The Geological Society of America. All rights reserved.

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Lericolais et al. lacustrine shelf deposits formed a significant basinward-prograding wedge system, interpreted as forced regression system tracts. On top of these prograding sequences, there is a set of sand dunes that delineates a wave-cut terrace-like feature around the isobath −100 m. The upper part of the last prograding sequence is incised by anastomosed channels that end in the Danube (Viteaz) canyon, which are also built on the lacustrine prograding wedge. Overlying this succession, there is a shelfwide unconformity visible in very high-resolution seismic-reflection profiles and present all over the shelf. A uniform drape of marine sediment above the unconformity is present all over the continental shelf with practically the same thickness over nearby elevations and depressions. This mud drape represents the last stage of the Black Sea water-level fluctuation and is set after the reconnection of this basin with the Mediterranean Sea.

INTRODUCTION The location of the Black Sea, between Europe and Asia, makes its water level dependent on Eurasian climatic fluctuations. This inland sea is a perfect present-day example of a marginal basin where connection changes dramatically with sea level (Ross, 1971, 1978; Ross et al., 1970; Ross and Degens, 1974; Ryan et al., 2003, 1997). The Black Sea is at present the world’s largest anoxic basin, making it an important modern analogue for past anoxic conditions, while during the last glacial period, it was a low-salinity oxygenated lake, isolated from the Mediterranean (Deuser, 1972, 1974; Lericolais et al., 2006b; Wall and Dale, 1974). The marine and lacustrine deposits at the Black Sea represent a valuable archive for the study of past climate changes. During the Quaternary glacial periods, a northern ice cap prevented major East European rivers flowing from north as they do today. During ensuing interglacial periods, these rivers were diverted to the south in the direction of the Black Sea and Caspian Sea receiving basins and consequently have increased the size of these drainage basins (Arkhipov et al., 1995). Therefore, unique conditions specific to the Black Sea were established while this water body became isolated from the global ocean. This isolation results in a sedimentary record without the hysteresis effect, which is the latent period needed by the global ocean to respond to the consequences of ice melting. During these isolation phases, the Black Sea was more sensitive to climate changes than the Caspian Sea is today. Arkhipov et al. (1995) and Chepalyga (1984) interpreted the Caspian Sea fluctuations opposed to those of the global ocean to have caused the possible connection between the Black Sea and the Caspian Sea through the Manych Strait (Fig. 1). When the Black Sea was isolated, both the lack of saltwater input and the increase of freshwater runoff from the rivers led to reduced salinity levels in the Black Sea. This process during the glacial periods, linked to water-level fluctuation, is measured in the fauna succession, which shows an abrupt change from saltwater to freshwater or brackish-water species. The initial hypothesis of a rapid saltwater flooding of the freshwater lake that was the Black Sea in the Late Glacial Maximum (LGM) was proposed in 1996 by Ryan et al. (1996, 1997). The flood hypothesis raised controversy and initiated refutation (Aksu et al., 2002a, 2002b, 1999; Görür

et al., 2001; Hiscott and Aksu, 2002; Hiscott et al., 2008, 2002; Yanko-Hombach et al., 2006), but recently also received support (Algan et al., 2007; Eriş et al., 2007, 2008; Gökaşan et al., 2005; Lericolais et al., 2007b, 2007c; Siddall et al., 2004). Nevertheless, most of each opposing view is supported by only a small amount of data in the Black Sea, and not all of the 420,000 km2 have been surveyed using modern scientific equipment and interpretation in light of modern ideas. Recently, the European Project ASSEMBLAGE (EVK3CT-2002-00090) provided geophysical and sedimentary data collected in the northwestern part of the Black Sea from the continental shelf and slope down to the deep-sea zone. This project focused on applying sequence stratigraphic models to seismic data recorded on the northwestern Black Sea shelf, in order to correlate the sequences interpreted using seismic stratigraphy methods to sea-level fluctuations. To achieve the project’s objectives, very high-resolution seismic data were acquired during the BlaSON cruises (1998 and 2002) using the research vessel Le Suroît and during the ASSEMBLAGE 1 (2004) cruise of the research vessel Le Marion Dufresne. During the first two cruises, paleoshorelines and sand ridges were identified, and a set of seismic data was acquired on these targets to support pseudo–three-dimensional (3-D) analyses. This, coupled with a multiproxy approach, emphasizes that the Black Sea water level is dependent on Eurasian climatic fluctuations. This sequence stratigraphy study was validated by dated samples obtained from long cores (up to 50 m long) providing a firm calibration of Black Sea water-level fluctuation since the LGM. These data show that the Black Sea experienced a contemporary rise in water level with the melting of the Fennoscandian ice sheet, followed by a drop of the water level from the Younger Dryas to the Preboreal. This recent lowstand is confirmed by the presence of the forced regression sequences, the wave-cut terrace, and the coastal dunes still preserved on the shelf, even after the Black Sea was rapidly invaded by Mediterranean/Marmara marine waters. PREVIOUS STUDIES Already in the seventies, Kuprin et al. (1974) and Shcherbakov et al. (1978) documented the lowstand shorelines of the Black

Assessment of Black Sea water-level fluctuations Sea. Numerous Soviet, Romanian, Bulgarian, and Turkish coring and echo-sounding surveys conducted in the western part of the Black Sea had previously identified a littoral zone near the shelf edge. Several cores cut by these studies penetrated an erosional surface. From the 1990s Romanian data, Popescu et al. (2004) identified the presence of ancient river valleys entrenching the shelf, especially in front of major canyons. Other workers confirmed the shoreline position with the recovery of sand, gravel, and freshwater molluscs typical of the coastal zone (Major et al., 2002b; Ryan et al., 2003). Ostrovskiy et al. (1977) published results on the stratigraphy and geochronology of Pleistocene marine terraces of the Black Sea, where extensive down-cutting of coastal river valleys was recognized as evidence of a major water-level drop of the ice-age Black Sea on the order of −110 m. A key limitation of this previous research is that no seismic-reflection profiles were published to document their findings, even though one can read that the former eastern country researchers have documented the exposed margin of the Black Sea lake, and that numerous piston and drill cores also confirmed the existence of an ancient coast. The first interna-

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tional publication related to the work done by the Soviet block on this topic was issued from the U.S.-Russia-Turkey expedition of 1993 led by professor Shimkus (Shimkus et al., 1997) with the objective to examine the impact of the Chernobyl contamination in the Black Sea. The results obtained from this joint survey allowed the mapping of the Dnieper River valleys in more detail with reflection profiling methods and explored the coastal deltas on the Ukrainian shelf. Later reflection profiling gave evidence of the same shelfwide erosion surface at different Black Sea locations, i.e., on the Romanian shelf (Lericolais et al., 2007b; Popescu et al., 2004), on the Bulgarian shelf (Dimitrov and Peychev, 2005; Dimitrov, 1982; Khrischev and Georgiev, 1991), and on the Turkish margin (Aksu et al., 2002b; Algan et al., 2002, 2007; Demirbag et al., 1999; Okyar and Ediger, 1999; Okyar et al., 1994). The general assumption about the Black Sea before the Ryan et al. (1997) hypothesis was that the lake’s surface had risen correlatively with the global sea level. This required a relatively early connection through the Bosporus Strait. Based on hydrologic considerations, Chepalyga (1984) and Kvasov and

Figure 1. General location of the studied area.

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Blazhchishin (1978) stipulated that outflow from the Black Sea through this strait had always been continuous, even at maximum lowstand conditions. For this to be the case, the outlet level of the Black Sea during glacial times would have to have been in concordance with the Black Sea lowstand shorelines. In opposition to the evidence discovered about the depth of the Bosporus sill, Chepalyga (1995) used an old suggestion to place the former Black Sea–Marmara connection in the Sakarya River valley. This idea was consistent with the work of Elmas (2003) who demonstrated that the late Cenozoic tectonics and stratigraphy of northwestern Anatolia allowed a connection between the Sakarya outlet to the eastern arm of the Gulf of Izmit. If such theory is viable, the Black Sea sill controlling its lake level would therefore have been the Dardanelles bedrock sill that is at −85 m (Ryan, 2007). Recently, a number of researchers have rejected the hypothesis of a deep Black Sea outlet (Bahr et al., 2005, 2006; Major et al., 2002b; Myers et al., 2003). Previously, Lane-Serff et al. (1997) had proposed that a deep outlet would have permitted a vigorous outflow of semibrackish waters from the Black Sea strong enough to keep out Mediterranean saltwater from entering the enclosed basin. However, their hydraulic models only prevented Mediterranean inflow until sea level rose 5 m above the sill. It is now admitted that the increase in salinity in the Sea of Marmara occurred at least 12,000 yr ago, as determined from the mollusc assemblage and stable isotopes (Çagatay et al., 2000; Sperling et al., 2003) and could have been even earlier (Popescu, 2003, 2004). Authors who are against a late connection of the Black Sea to the Mediterranean sea are now publishing evidence of a late salinization of the Black Sea obtained from their studies conducted on cores recovered on the Black Sea Turkish shelf, e.g., Hiscott et al. (2007) explains that the Ostracoda of Caspian affinity indicate ~5‰ salinity until ca. 7500 yr B.P. Dinocysts and foraminifera confirm a low but rising salinity no later than ca. 8600 yr B.P., and a first major pulse of marine waters was recorded at around 8460 yr B.P. by Marret et al. (2009). Hence, they confirm the previous observations published by Ryan et al. (1997, 2003) and Major et al. (2002a, 2002b, 2006). These results propose that the first marine signal in the Black Sea is recorded between 9000–8000 yr B.P. At that time, the Mediterranean sea level was around −30 m (Lambeck and Bard, 2000; Lambeck et al., 2002) or even less, as the model predictions from the northern coast of Israel indicate sea level at about –13.5 ± 1 m for ca. 8000 yr B.P., whereas observation places it between –14.5 m and –16.5 m (Lambeck et al., 2007; Sivan, 2003). If the Black Sea outflow through a deep connection was truly so vigorous and persistent, it remains to be explained how this outflow could have permitted the early and sustained salinization of the Sea of Marmara at the downstream end of the water cascade. On the other hand, since the observation of post-LGM lowstand shorelines characterized by wave-cut terraces in different areas of the Black Sea, i.e., at –110 m off Ukraine (Ryan et al., 1997), −100 m on the Romanian shelf (Lericolais et al., 2003, 2007b, 2007c), −122 m for the Bulgarian shelf (Dimitrov, 1982),

and −155 m off Sinop where the shelf is really narrow (Ballard et al., 2000), it is necessary to consider a shallow outlet for the Bosporus with interrupted outflow. Even if the subsidence effect since the LGM is negligible (Wong et al., 2005), some consideration has to be taken for the tectonic effect, especially at the foot of major mountain chains such as the Carpathians, the Balkans, the Caucasus, or Anatolia. This effect explains why some 20 m of difference exists for the topset location of the LGM lowstand wedges in different Black Sea areas. Major et al. (2006), quoting their former work on strontium isotopes published in 2002 (Major et al., 2002b), established that the lake level would have been controlled mainly by the balance of evaporation versus inflow from rivers and rainfall, even though intervals of enclosure of the Black Sea may have been of relatively short duration. These authors also confirm that the Black Sea was an enclosed semibrackish lake during these periods. Lake-level fluctuations might also account for the observed repetition of “cut and fill” in the sediments of the river valleys that cross the shelf (Heller et al., 2001; Koss et al., 1994; Lericolais et al., 2001; Newell, 2001; Popescu et al., 2004; Ryan et al., 2003; Talling, 1998; Zaitlin et al., 1994), as well as the presence of wave-cut terraces on the edges of the shelf (Shimkus et al., 1980). The presence of authigenic aragonite layers correlative with the onset of the sapropel deposit (Giunta et al., 2007) can be correlated to a response to climatic change (Lamb, 2001) despite the hydrothermal influence and calcite precipitation (Peckmann, 2001). A detrital/biogenic source has also been interpreted by Reitz and de Lange (2006) as a possible mechanism for the major part of the aragonite enrichments found in sapropel sediments. Possibly, offshore-directed surface-water flows related to wind stress and/ or enhanced runoff (consistent with Mediterranean flooding and enhanced precipitation) during sapropel deposition may have assisted in the transport of near-coastal aragonitic organisms to more coast-remote areas. Recent studies carried out in the Black Sea confirm that authigenic calcite precipitation of calcareous mud appears following the deglacial meltwater delivery (Bahr et al., 2005, 2006; Major et al., 2002b; Ryan et al., 2003) and can be interpreted as a result of water evaporation (Giunta et al., 2007). SYNTHESIS OF RESULTS OBTAINED IN THE FRAME OF THE ASSEMBLAGE PROJECT Recently, an assessment of the northwestern part of the Black Sea sedimentary systems from the continental shelf and slope down to the deep-sea zone was provided by the ASSEMBLAGE European Project. Here, we summarize the results of this project, obtained from geophysical data and core analyses. These results provide a solid record of the Black Sea Last Glacial Maximum (LGM) water-level fluctuations and shed new light on the controversy concerning the Black Sea water-level fluctuation since the Last Glacial Maximum. The ASSEMBLAGE project attempted to assess the last sealevel rise in the Black Sea and provide scenarios quantifying the processes governing the transition of the Black Sea system from

Assessment of Black Sea water-level fluctuations a low-salinity lake to a marine state while addressing the variability in this system. Six major observations are used to reconstruct the Black Sea water-level fluctuations since the LGM. 1. The first observation is the existence of a LGM lowstand wedge at the shelf edge offshore Romania, Bulgaria, and Turkey. This observation is completed with the evidence of a second small lowstand wedge dated from 11,000 yr B.P. to 8000 yr B.P. from −100 to −120 m of water depth identified during ASSEMBLAGE cruises on the outer shelf of Romania and Bulgaria and described on the Turkish shelf by Algan et al. (2002). This wedge is associated with the recovery of strata immediately below an observed unconformity consisting of dense, low-water-content mud containing desiccation cracks, plant roots, and sand lenses rich in freshwater molluscs (Dreissena rostriformis) with both valves still together. 2. The second observation is deduced from results providing information on the construction of the Danube delta/prodelta, showing that a former prodelta was built up at −40 m after the post-LGM meltwater pulses. 3. The third observation comes from mapping of meandering river channels capped by a regional unconformity and extending seaward across the Romanian shelf to the vicinity of the –100 m isobath. 4. The fourth observation is the presence of submerged shorelines with wave-cut terraces and coastal dunes, or delta mouth bars at depths between –80 to –100 m, below the Holocene Bosporus and Dardanelles Strait outlet sill to the global ocean. 5. The fifth observation to be underlined is that, on the western part of the Black Sea continental shelf, a shelfwide ravinement surface is visible in very high-resolution seismic-reflection profiles. 6. The sixth observation useable for the understanding of the last water-level fluctuation of the Black Sea is the presence of a uniform drape of sediment beginning at the same time above the unconformity with practically the same thickness over nearby elevations and depressions and with no visible indication of coastal-directed onlap across the outer and middle shelf, except in the vicinity of the Danube Delta, where this mud drape is overlapped by recent Danube sediments.

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digitally recorded. The navigation profiles presented here are displayed in Figure 2. The very high-resolution seismic sources were hull-mounted Chirp sonar systems with frequencies ranging between 1.5 and 7 kHz. Their vertical resolution is less than 1 m, with penetration reaching 500 ms in some deep areas where the sediment cover is constituted by soft sediment. On the shelf, the presence of gasbearing sediments masking the information beneath ~20 m depth below seafloor (bsf) decreases the amount of usable data. We also present results obtained from cores. The sediment cores were recovered using a Kullenberg piston corer; a conventional one used during BlaSON surveys, and Calypso type one developed for IPEV (French Institute for South Polar Seas) with a long tube (up to 60 m) system. Each of the core sections recovered were cut horizontally into two pieces and scanned to get an image before analyzing the samples. The general properties of the sediments were measured by the Multi System Track (MST) to get P-wave velocity and amplitude, density, impedance, and magnetic susceptibility values. All cores were packaged at 4 °C, and sampling was done at the IFREMER Brest laboratory. Dating was conducted on samples at various distances and various depths from the coast to reveal any possible bias in ages due to coastal or current influence. The Poznań Radiocarbon Laboratory in Poland performed 14C dating.

Methods of Data Acquisition The data were acquired during surveys realized in the framework of two main projects; (1) BlaSON: a French-Romanian bilateral project for which two surveys coordinated by IFREMER (Institut Francais de Recherche pour l’Exploitation de la Mer) were carried out on board the French RV Le Suroît in 1998 and 2002, and (2) ASSEMBLAGE: an FP5 European project for which two surveys were carried out on board the French RV Le Marion Dufresne in 2004 and the Romanian RV Mare Nigrum in 2005. For all these surveys, a differential global positioning system (GPS) was deployed for accurate (~1 m) positioning, and every vessel was equipped with swath bathymetry systems. Very high-resolution seismic lines were shot simultaneously using a Chirp sonar system. All data acquisition was synchronized and

Figure 2. Bathymetry of the semi-enclosed Black Sea basin and BlaSON and ASSEMBLAGE survey route locations.

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Before the evidence of wave-cut terraces in the Black Sea, no reliable sea-level markers were described to allow a good sea-level reconstruction (Giosan et al., 2006; Pirazzoli, 1991). Moreover, the lack of radiocarbon ages on in situ materials and the difficulty in calibrating radiocarbon ages in a setting with variable reservoir ages (Giosan, 2007) led to ongoing discussion about the Black Sea water-level fluctuations. At least for the more recent period, Siani et al. (2000) have proposed a reservoir age of 415 ± 90 yr B.P. for the Black Sea, based on six samples from the Black Sea, the Sea of Marmara, and in the vicinity of the Bosporus. One of the reasons for this is that a reservoir age of ~1280 yr was deduced from the occurrence of the Santorini Minoan ash in the Unit II of Jones and Gagnon (1994) from several south Black Sea cores recovered between 400 and 700 m below present sea level (Guichard et al., 1993). If we include documentation from Jones and Gagnon (1994), Bahr et al. (2005), and Kwiecien et al. (2008), then reservoir ages extending from 0 to 1280 yr. For the lacustrine period, no measurement has been proposed for the past. If no terrestrial influence exists, then age and residence time of deep waters will be the main factors. Östlund (1974) calculated ages of deep waters between 1470 yr (at ~600 m) and more than 2000 yr (when deeper than 1400 m). If a low stratification occurs in poorly salted water, then the residence time would be equal or less than 935 yr. The question of which reservoir age should be given to old water of the “Black Lake,” depending on depth in the water column, is still a matter of debate, and it is the reason why we still use uncalibrated and uncorrected 14C ages throughout this study for our obtained dates. The seismic profile and core locations are displayed on Figure 3. Core location, length, and water depth where samples were recovered are presented in Table 1. First Observation: The Lowstand Wedges (LGM and Preboreal) The seismic line Chirp B2CH96 (AB on Fig. 3) was shot during the BlaSON2 survey off Romania in front of the Danube delta (Fig. 4). At around 150 m water depth, this line displays prograding parallel but undulating reflectors characterizing a seismic unit LSW1 (Fig. 5). These reflectors toplap at the top of LSW1. Above the erosional truncation, unit LSW2 is located on the slope part of this dip line. This unit presents reflectors beveling the LSW1 slope slightly to the northwest. Throughout the seismic line, there is a thin unit that corresponds to the mud drape known to be present all over the western Black Sea (Lericolais et al., 2007b; Major et al., 2002b; Ryan et al., 1997, 2003). Age control of these two lowstand wedges is given by the dates obtained from core MD04-2771 and presented in Table 2. Dating of the seismic units interpreted as the LSW1 and LSW2 was possible, and older dates reach back to 29,450 ± 320 14C yr B.P. for unit LSW1 at 11.90 m on core MD04-2771. This date was obtained on organic matter, but because a Dreissena shell sampled at 2.18 m

in the same core returned a date of 24,980 ± 160 14C yr B.P., we can be confident in attributing the deposition of LSW1 to the Last Glacial period. A second lowstand is evidenced by seismic sequence LSW2. This lowstand wedge has a shape characteristic of a low-energy wedge. Core MD04-2771 confirms that this lowstand wedge started to be deposited around 12,180 ± 60 14C yr B.P. On the southwestern part of the Black Sea, another Chirp profile (B2CH56) displays more precisely the two successive lowstand wedges LSW1 and LSW2 (Fig. 5). Age control of these lowstand deposits was obtained from core MD04-2752 (Table 3; Fig. 6) dating Dreissena shells. Here again, the LSW1 is correlative to LGM time. It is very clear that LSW2 is a lowstand wedge deposited between 12,010 ± 50 14C yr B.P. and 8130 ± 50 14C yr B.P., showing that the Black Sea encountered a second lowstand after the LGM lowstand. Second Observation: The Post-LGM (Ante–Younger Dryas) Danube Prodelta The Danube prodelta is located at the coastal part of the Danube delta and can be seen on line B2CH96, section AB (Fig. 3). On the Chirp seismic profile, an erosion surface interpreted as a ravinement surface R1 is identified (Fig. 7). Above this, a prograding wedge U.S.2 is well delimited on the Chirp seismic data; this wedge corresponds to a former prodelta lobe. Another prodelta lobe corresponding to seismic unit U.S.3 presents reflectors onlapping on the previous unit U.S.2. The geometric relationship between these prodelta lobes would have been best imaged on shore-parallel profiles showing in detail the lap-out patterns of seismic reflectors, such as those shown for the Po River prodelta by Correggiari et al. (2005a, 2005b). Units U.S.2 and U.S.3 (Fig. 7) represent the sites of deposition and progradation at the distal part of the previous Danube River outlets and channels. These prodelta lobes represent part of the main depocenters that extend offshore, being a considerable portion of the prodelta deposit (Correggiari et al., 2005a, 2005b). As seen on Figure 4, these units are the highest part of the dip seismic line and are restricted to the prodelta area. Our data set is not dense enough to be able to decipher prodelta autocyclic processes from external forcing. Nevertheless, their position and nature are in accordance with our interpretation obtained from the comparison of the seismic data interpreted all along the profile and the core results. Above these prodelta lobes, seismic unit U.S.4 (Fig. 7) is prograding also, but in a more gentle shape. This unit can also be interpreted as a prodelta lobe, but ages obtained on core MD042774 return an average age of ca. 9500 14C yr B.P. (see Table 4). U.S.4 is contemporaneous to the onset of the Preboreal regression responsible of the deposit of LSW2 at the shelf edge as presented in the previous paragraph. Recent studies (Giosan et al., 2009) confirm that the Danube was building a ramp delta lobe at 8860 ± 45 14C yr B.P. (ages obtained from Dreissena polymorpha). From the morphology of

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Figure 3. Western Black Sea shelf presenting the location map of the seismic line B2CH96 (A–B) and B2CH56 (C– D) and of the cores MD04-2752, MD042771, MD04-2772, MD04-2773, and MD04-2774, along with geomorphologic interpretation issued from previous work (Lericolais et al., 2007c; Popescu et al., 2001, 2004). DA—dune area, DD—Danube delta, PDR1 and PDR2—paleo–Danube River 1 and 2, PCL—paleocoastline, VC—Viteaz Canyon, DSF—Danube deep-sea fan, BSF—Bosporus shallow fan delta.

TABLE 1. CORES PRESENTED IN THIS STUDY WITH THEIR POSITION, LENGTH OF RECOVERY, AND WATER DEPTH AT LOCATION Latitude (°N) Longitude (°E) Core Core length Water depth (m) (m) MD04-2752 41°56.76 28°36.56 24.50 169 MD04-2771 44°16.32 30°54.24 12.38 168 MD04-2772 44°18.07 30°51.56 7.51 106 MD04-2773 44°37.96 30°20.61 3.63 68 MD04-2774 44°57.47 29°50.12 7.30 30

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Figure 4. Line B2CH96 with core location (vertical exaggeration is ~200). DPD—Danube prodelta, PCL—paleocoastline, LSW—lowstand wedge, f—faults; twt—two-way traveltime.

Figure 5. Chirp profile B2CH96: Distal part of Figure 4. Lowstand wedge 1 (LSW1) and LSW2 can be distinguished. twt—two-way traveltime.

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TABLE 2. DATES OBTAINED FOR THE LOWSTAND WEDGES FROM CORE MD04-2771 AND TYPE OF SAMPLES DATED (MOLLUSC OR ORGANIC MATTER) Calibrated age* Water depth Core length Depth in core Age Sample Unit Core 14 (yr) (m) (m) (m) ( C yr B.P.) 0.34 12,180 ± 60 13,650 ± 120 Dreissena LSW2 MD04-2771 168 12.38 2.18 24,980 ± 160 28,840 ± 180 Dreissena LSW1 – 11.90 29,450 ± 320 Organic matter LSW1 *Calibrated ages are here as indicator and were obtained using the Radiocarbon Calibration Program Calib5 (Stuiver et al., 1998) with 400 yr for reservoir correction.

TABLE 3. DATES OBTAINED FOR THE LOWSTAND WEDGES FROM CORE MD04-2752 AND TYPE OF SAMPLES DATED (MOLLUSC OR ORGANIC MATTER) Calibrated age* Water depth Core length Depth in core Age Sample Unit Core 14 (yr) (m) (m) (m) ( C yr B.P.) 12.20 8130 ± 50 8605 ± 110 Organic matter LSW2 MD04-2752 169 24.50 12.30 12,010 ± 50 13,430 ± 100 Dreissena LSW2 19.85 25,020 ± 180 28,730 ± 300 Dreissena LSW1 *Calibrated ages are here as indicator and were obtained using the Radiocarbon Calibration Program Calib5 (Stuiver et al., 1998) with 400 yr for reservoir correction and Cariacco data for age >24,000 yr.

the lacustrine-marine contact, Giosan et al. (2009) supposed that the Black Sea lake level at that time was around 30 mbsl. Third Observation: Meandering River Channels Preserved on the Black Sea Shelf The third observation is deduced from previous Romanian surveys carried out by the GeoEcoMar Institute, where several

recent paleoriver channels incising the continental shelf down to −90 m water depth (Popescu et al., 2004) were identified (PDR1 and PDR2 on Fig. 3). These paleochannels are completely filled by sediments and are no longer visible in the bathymetry. These erosive features reach 400–1500 m in width and 20–30 m in depth. They present conventional asymmetry on some cross sections (Fig. 8) and seem to have been beveled by a subsequent phase of erosion. Here also these paleochannels are sealed by the

Figure 6. Chirp profile B2CH56: Distal part of segment line C–D on Figure 3 with core MD04-2752 location. Lowstand wedge 1 (LSW1) and LSW2 can be distinguished. Dated core samples are a = 8130 ± 50 14C yr B.P., b = 12,010 ± 50 14C yr B.P., c = 25,020 ± 180 14C yr B.P. (cf. Table 3). twt—two-way traveltime.

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Figure 7. Prodelta section of line B2CH96 with location of core MD04-2774: R1 is a ravinement surface; U.S.2 is a prograding wedge interpreted as a former prodelta lobe; U.S.3 is a retrograding sequence on U.S.2 and can be a second prodelta lobe; U.S.4 is a prograding sequence to be correlated to the forced regression described at −100 m by Lericolais et al. (2006a); and U.S.8 is the present-day prodelta deposit.

mud drape described earlier (Lericolais et al., 2007a; Major et al., 2002b; Popescu et al., 2004; Ryan et al., 2003). For Popescu et al. (2004), the stratigraphic position of these incisions lying directly under the discontinuity at the base of the Holocene strongly suggested that they formed during the last lowstand. The cartography of these buried channels shows that they are concentrated around two main directions. This distribution leads to their interpretation as anastomosed fluvial systems corresponding to two distinct drainage systems (Fig. 3). These would correspond to former paleo–Danube River flooding on the shelf to the outer shelf, where they apparently split into several arms, similar to a fluvial deltaic structure comparable in size to the modern Danube

delta, that lie close to the Danube Canyon (Popescu et al., 2004), also named Viteaz Canyon (VC on Fig. 3). The channels extend right to the paleoshoreline and pass under the belt of coastal sand ridges and depressions. Consequently, the regression that exposed the shelf surface into which the river channels were cut was followed by a transgression that led to the filling of the channels and then to another regression that deflated the channel fills and reexposed the entire region to coastal dune and pan development. The argumentation about the origin of the coastal features at ~−100 m has been presented in Lericolais et al. (2007b). Core MD04-2773 was recovered at one incised valley section (Fig. 8). The core (Fig. 9) got through the marine drape and

TABLE 4. DATES OBTAINED ON THE CORE MD04-2774 AND TYPE OF SAMPLES DATED (MOLLUSC OR ORGANIC MATTER) Calibrated age* Sample Unit Core Water depth Core length Depth in core Age 14 (yr) (m) (m) (m) ( C yr B.P.) 5.43 9030 ± 50 9720 ± 140 Pisidium US4 MD04-2774 30 7.3 6.91 9570 ± 50 10,440 ± 90 Pisidium US4 *Calibrated ages are here as indicator and were obtained using the Radiocarbon Calibration Program Calib5 (Stuiver et al., 1998) with 400 yr for reservoir correction.

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Figure 8. Incised valley section of line B2CH96 with location of core MD04-2773. TWT—two-way traveltime.

passed the Dreissena hash layer (Major et al., 2002b) described later herein. Even if the mollusc Pisidium sampled at 103 cm in the core (Table 5) is at the limit of the hash layer, it shows that this mollusc is of the same genus as the one collected in core MD042774. It is a genus of very small or minute freshwater clams known as pea clams, aquatic bivalve molluscs in the family Sphaeriidae. This confirms that the infilling of the valleys was still active at more than 60 m before the onset of the marine drape. Fourth Observation: Presence of Submerged Shorelines The submerged shorelines characterized by the presence of a wave-cut terrace at depths between –80 and –100 m are the key elements of the fourth observation. At the top of this coastal feature recognized on the Romanian shelf, there is a set of coastal dunes or delta mouth bars described by Lericolais et al. (2007a, 2007b). Analysis of the very high-resolution seismic data in pseudo–3-D mode (Lericolais et al., 2009) demonstrates that the lacustrine shelf deposits form an important basinward-prograding wedge system interpreted as a forced regression system tract eroded at the distal part by a wave-cut terrace (see figs. 4 and 5 in Lericolais et al., 2009). On line B2CH96, located north of the dune

field studied area, the wave-cut terrace is also visible (Fig. 10). On top of the prograding units (FR on Fig. 10), there is a set of sand dunes that delineates a berm-like feature around the −100 m isobath (WCT on Fig. 10), similar to the ones described by Ryan et al. (2003), Popescu et al. (2004), and Lericolais et al. (2007b). Analyses of cores retrieved from the dune field area demonstrate that the prograding wedges are lacustrine in origin and document a low water level characterized by forced regression– like reflectors mapped from the pseudo–3-D seismic data (Lericolais et al., 2009). Here, too, the hinge point corresponds to the wave-erosion surface mapped around the −100 m isobath. The ages returned by the core analysis range between 11,000 and 8000 14C yr B.P., with the formation of dunes being around 8500 14C yr B.P. The prograding reflectors deepen seaward and are truncated by an erosional surface described as the wave-cut terrace. On the Chirp profile, it is clearly seen that all the area is covered by a drape of less of 1 m thick (see “Sixth Observation: Uniform Drape above the Unconformity”), confirming that the dune system is not active any more. Everywhere across the midand outer shelf, the ridges, mounds, and depressions are draped by this thin layer of sediment with a remarkably uniform thickness of no more than a meter (Fig. 11).

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Lericolais et al. where the present-day washed wave zone is marked by debris of whitened mollusc shells. Such facts are in favor of a rapid transgression in the Black Sea and are in agreement with the previous works published by Khrischev and Georgiev (1991) and Lericolais et al. (2004, 2007b). Actually, Khrischev and Georgiev (1991) attributed “fast rising” water level to the transition from lacustrine to marine conditions. For them, this change corresponds to a stratigraphic break (“washout”) in the cores that interrupts the lacustrine calcite precipitation and is followed by terrigenous mud with marine molluscs. They reported this “washout” in more than 100 cores. This same transition was described for the BlaSON and ASSEMBLAGE cores, where the transition was interpreted as either a ravinement surface or an erosion surface (Lericolais et al., 2007b, 2007c; Major et al., 2002b). Algan et al. (2007, p. 621) also made the observation of dense, dry mud below the erosional unconformity on the Thrace margin in cores from the shelf edge. These authors noted “a marked contact” between a 2-cm-thick shell-enriched layer and a “stiff clay deposit with low water content at the base of these cores.” The 14C age of the shells (Dreissena sp.) is 8590 ± 145 yr B.P., comparable to the age of the shell material constituting the “hash layer.” Algan et al. (2007, p. 623) considered that “the lithological characteristic of this core indicates that the deposition starts with high-energy condition over the stiff eroded substrate at about −100m, and continued with low-energy, suggesting a rapid deepening of a shallow environment.” Sixth Observation: Uniform Drape above the Unconformity

Figure 9. Photo of the core MD04-2773 sections. I—Modiulus ecozone; II—Mytillus ecozone, III—Dreissena ecozone. I and II are marine indicators, while III is from semibrackish state (Giunta et al., 2007). a—Pisidium at 103 cm in the core aged 7890 ± 50 14C yr B.P.

Fifth Observation: Ravinement Surface On the western part of the Black Sea continental shelf, a shelfwide ravinement surface is always present and can be recognized both on very high-resolution seismic-reflection profiles and in all the collected cores. In the cores, this surface corresponds to the described “hash layer” of Major et al. (Major et al., 2002b). This “hash layer” is composed of debris of whitened Dreissena. This corresponds to the surf zone as it is shown on Figure 12

Along the Black Sea margin, Wong et al. (2005), Algan et al. (2002), Ryan (2007), Ryan et al. (2003), Major et al. (2002b), and Lericolais et al. (2007b, 2007c) already described the presence of a uniform mud drape deposited above the unconformity. This mud drape layer was sampled during BlaSON and ASSEMBLAGE and corresponds in cores to the layer of terrigenous mud containing marine molluscs such as Mytilus galloprovincialis and Mytilus edulis, Cerastoderma edule, and Cardium edule (Giunta et al., 2007). This lithologic and biostratigraphic interval on the shelf corresponds to units 1 and 2 in basin sediments as defined by Ross et al. (1970). This uniform mud drape is clearly seen on the highresolution seismic profiles obtained by the Chirp sonar system and is displayed for instance on Figures 8 and 11. Its thickness, when calculated from acoustic travel time to meters, corresponds

TABLE 5. DATES OBTAINED ON CORE MD04-2773 AND TYPE OF SAMPLES DATED (MOLLUSC OR ORGANIC MATTER) Calibrated age* Sample Unit Water depth Core length Depth in core Age 14 (yr) (m) (m) (m) ( C yr B.P.) MD04-2773 68 3.63 1.03 7890 ± 50 8350 ± 60 Pisidium Incised valley *Calibrated age are here as indicator and were obtained using the Radiocarbon Calibration Program Calib5 (Stuiver et al., 1998) with 400 yr for reservoir correction. Core

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Figure 10. Forced regression (FR) progradation limited basinward by the wave-cut terrace (WCT) visible on section of line B2CH96 with location of core MD04-2773 (note the prograding reflectors inside the FR wedge). TWT—two-way traveltime.

in cores to the layer of terrigenous mud containing marine molluscs. Similar to the Romanian continental shelf, this layer has also been found above the unconformity on other Black Sea margins (Algan et al., 2002, 2007). Initial deposition of this uniform drape of sediment started at the same time above the unconformity and has practically the same thickness over nearby elevations and depressions, and it presents no visible indication of coastal-directed onlap across the outer and middle shelf. Such a layer deposited over the “hash

layer” ravinement surface, composed of in situ mussel molluscs at the bottom of this infra-meter layer, is characteristic of a rapid change. The size and disposition of the Mytilus edulis found in the cores are in accordance with the natural biotope of such a species. While the highest biomass is in general recorded at water depths ranging between 5 and 30 m, being lower at deeper depths, and living in niche beyond 40 m (Stea et al., 1994; Westerbom et al., 2002), this is not the case in our cores, where they are abundant everywhere. Such an observation is also an argument in favor of

Figure 11. Close-up of Figure 10 Chirp profile B2CH96 showing the seismic signal of the mud drape. TWT— two-way traveltime.

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a rapid sea-level rise, making it hard for these mussels to survive the transgression. SYNTHESIS: WATER-LEVEL FLUCTUATION OF THE BLACK SEA SINCE THE LAST GLACIAL EXTREME The synthesis presented here corresponds to an essay on the assessment of the last sea-level rise in the Black Sea and supports a scenario quantifying the processes governing the transition of Black Sea system from a semi-freshwater lake to a marine state. This work addresses the important postglacial variability in the Black Sea system, as the transition of this system from a semifreshwater lake to a marine environment was perhaps one of the most dramatic late Quaternary environmental events in the world. Back to the Last Glacial Maximum, 21,000 yr ago, the Black Sea was probably a giant freshwater to semi–brackish-water lake, as proposed by Arkhangelskiy and Strakhov (1938), or at least a brackish enclosed basin. Its water level stood more than 120 m below than today’s level. During ASSEMBLAGE project, analysis of high-resolution seismic-reflection profiles, Chirp and side-scan data together with piston core analyses from surveys taken on the Danube fan, and on the Black Sea shelf, provided new insights into the recent sedimentation processes in the deep northwestern Black Sea. The deep-sea fan studies (Popescu et al., 2001) demonstrate that the last channel-levee system on the Danube fan developed during the Neoeuxinian lowstand (stage 2) in a semi-freshwater basin with a water level ~120 m lower than today. Sediments supplied by the Danube were transported to the deep basin through the Viteaz canyon (Popescu et al., 2004). Functioning of the deep-sea fan is a good indicator of lowstand periods (Popescu et al., 2001; Winguth et al., 2000; Wong et al., 1997).

Figure 12. The present-day wave action zone showing debris of coquinas at the berm. This hash layer is the modern equivalent of the one described in the Black Sea core and having an average age of 8600 14C yr B.P.

Because the Black Sea was in a very close vicinity to the Scandinavian-Russian ice cap, the supply of the melting water from the glaciers into the Black Sea through the major drainage system constituted by big European rivers (Danube, Dnieper, Dniester, and Bug) was recorded by a brownish layers described in cores (Bahr et al., 2005; Major et al., 2002b). The water volume brought to the Black Sea after the meltwater pulse 1A (MWP1A) at ca. 12,500 14C yr B.P. (14,500 yr cal. B.P.) (Bard et al., 1990) was sufficient to raise the water level between −40 m and −20 m, where the Dreissena layers were deposited. The −40 m upper limit is interpreted from our records and especially deduced from the construction of the Danube prodelta (Lericolais et al., 2009), which are not exhaustive, and the −20 m limit is certified by Yanko (1990). This last value for the transgression upper limit would have brought the level of the Black Sea even higher to the Bosporus sill, and possible inflow of marine species like Mediterranean dynoflagellate populations can be envisaged (Popescu, 2004). Nevertheless, the rise in the Black Sea water level, which stayed between freshwater to brackish conditions, stopped the deep-sea fan sedimentation. Palynological studies conducted on BlaSON cores (Popescu, 2004) show that from the Bølling-Allerød to the Younger Dryas, a cool and drier climate prevailed. Northeastern rivers converged to the North Sea and to the Baltic Ice Lake (Jensen et al., 1999), providing reduced river input to the Black Sea and resulting in a receding shoreline. These observations are consistent with some evaporative drawdown of the Black Sea and are correlated to the evidence of an authigenic aragonite layer present in all the cores studied (Giunta et al., 2007; Strechie et al., 2002). This drawdown is also confirmed by the determination of the forced regression– like reflectors recognized either on the dune field mosaics (Lericolais et al., 2009) or on the B2CH96 transect profile and dated to this period. This lowered sea level in the Black Sea persisted afterward. The post–Bølling-Allerød climatic event favored the lowering of the Black Sea water level, and the presence of the coastal sand dunes and wave-cut terraces confirms this lowstand. This had already been observed by several Russian authors who considered a sea-level lowstand at about −90 m depth. Their observations were based on the location of offshore sand ridges described at the shelf edge south of Crimea. The anastomosed buried fluvial channels described by Popescu et al. (2004) that suddenly disappear below −90 m depth and a unique wave-cut terrace on the outer shelf, with an upper surface varying between −95 and −100 m, are therefore consistent with a major lowstand level situated somewhere around −100 m depth. Around the Viteaz Canyon, the paleocoastline was forming a wide gulf into which two rivers flowed (Fig. 3). Previous studies have already proposed a depth of −105 m for this lowstand, according to a regional erosional truncation recognized on the southern coast of the Black Sea (Demirbag et al., 1999; Görür et al., 2001), but also based on a terrace on the northern shelf edge (Major et al., 2002b). On the Romanian shelf, preservation of the sand dunes and buried small incised valleys are to be linked with a rapid transgression where the ravinement processes related to the

Assessment of Black Sea water-level fluctuations water-level rise had no time to erode sufficiently the sea bottom (Benan and Kocurek, 2000; Lericolais et al., 2004). Circa 7500 14C yr B.P., the surface waters of the Black Sea suddenly attained present-day conditions owing to an abrupt flooding of the Black Sea by Mediterranean waters, as shown by dinoflagellate cyst records (Popescu, 2004). This can also be related with the beginning of widespread and synchronous sapropel deposition across slope and basin floor. At 7160 14C yr B.P., Popescu (2004) demonstrated a sudden (<760 yr according the resolution of their data) inflow of a very large volume of marine Mediterranean waters, causing an abrupt increase in salinity that attained the present-day euxinic values. This inflow of marine waters is confirmed by the abrupt replacement of freshwater to brackish species by marine species. Furthermore, the model developed by Siddall et al. (2004) shows that ~60,000 m3 of water per second must have flowed into the Black Sea basin after the sill broke, and it would have taken 33 yr to equalize water levels in the Black Sea and the Sea of Marmara. Such a sudden flood would have preserved lowstand marks on the Black Sea northwestern shelf. From part of this synthesis and based on the pseudo–3-D geometry of the seismic data interpreted by Lericolais et al. (2009), the water-level fluctuation diagram proposed here (Fig. 13) fits these synthesized observations. CONCLUSIONS This synthesis, based on data collected in the Black Sea from a 10 yr project, provides a solid record of water-level fluc-

Figure 13. Water-level fluctuation in the Black Sea since the Last Glacial Maximum (LGM). MWP1—meltwater pulse 1A; B/A— Bølling-Allerød; YD—Younger Dryas; PB—Preboreal. SI to SIX are the sequences interpreted and dated from the Romanian Black Sea shelf. (SI dates are: 23,630 ± 180; 24,980 ± 200, and 26,630 ± 230 14C yr B.P. SIII date is: 10,100 ± 50 14C yr B.P. SII, SIV to SVI revealed no date. SVII dates are: 11,040 ± 50, 10,930 ± 50, and 11,090 ± 50 14C yr B.P. SVIII dates are: 8760 ± 40 and 8600 ± 50 14C yr B.P. SIX date is: 8620 ± 50 14C yr B.P.) Figure is modified from Lericolais et al. (2009).

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tuation during the Last Glacial Maximum in the Black Sea. The starting point of this synthesis is based on the evidence at the shelf edge in Romania, Bulgaria, and Turkey of a Last Glacial Maximum lowstand wedge. From the increase of water provided to the Black Sea by the melting of the ice after 18,000 yr B.P. and drained by the largest European rivers (Danube, Dnieper, Dniester), a Danube prodelta was built under ~40 m of water depth, corresponding to the ensuing transgressive system. Subsequently, the Black Sea lacustrine shelf deposits formed a significant basinward-prograding wedge system, interpreted as forced regression system tracts. On top of these prograding sequences, set of sand dunes delineates a wave-cut terrace-like feature around the isobath −100 m. These coastal features as well as the incised anastomosed channel system were preserved on the shelf because the final transgression was fast enough to preserve them. A uniform drape of marine sediment above the unconformity is present all over the continental shelf with practically the same thickness over nearby elevations and depressions. This mud drape represents the last stage of the Black Sea water-level fluctuation and is set after the reconnection of this basin with the Mediterranean Sea. From such behavior, it seems that the sedimentary sequences in the Black Sea were strongly affected by sea-level changes driven by global glaciations and deglaciations. The level of the Black Sea, to a certain extent, was controlled more by the regional climate than by global eustatic changes. The transition of the Black Sea system from a lacustrine to a marine environment is perhaps one of the best records of climate change on the European continent. Six major observations have documented this Black Sea behavior: (1) existence of two lowstand wedges, one dated from the LGM and covered by a second one dated from 11,000 yr B.P. to 8000 yr B.P. and located at a water depth ranging from −100 to −120 m; (2) a Danube prodelta built up at −40 m after the post-LGM meltwater pulses; (3) a set of meandering river channels capped by a regional unconformity and extending seaward across the Romanian shelf to the vicinity of the –100 m isobath; (4) evidence of submerged shorelines with wave-cut terraces and coastal dunes, or delta mouth bars at depths between –80 to –100 m, below Holocene Bosporus and Dardanelles Strait outlet sill to the global ocean; (5) observation on the western part of the Black Sea continental shelf of a shelfwide ravinement surface visible in very high-resolution seismicreflection profiles; and (6) the presence of a uniform drape of sediment beginning at the same time above the unconformity with practically the same thickness over nearby elevations and depressions and with no visible indication of coastal-directed onlap across the outer and middle shelf. ACKNOWLEDGMENTS Our research was supported by the French Ministry of Foreign Affairs within the framework of a bilateral collaboration between France and Romania, and prolonged by a European project of the 5th program called ASSEMBLAGE (EVK3CT-2002-00090). Special acknowledgment goes to Nicolae

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Panin, who started and supported the project. We also thank Eliane Le Drezen and Alain Normand for their work in multibeam data processing, Hervé Nouzé for seismic processing, the crew of the vessels Le Suroît (Institut Francais de Recherche pour l’Exploitation de la Mer) and Le Marion Dufresne (Institut Paul Emile Victor). Special thanks are due to Yvon Balut. Special “remerciement” goes to Oya Algan, Mike Sweet, and Michel Fontugne, who edited this manuscript, for their helpful comments. REFERENCES CITED Aksu, A.E., Hiscott, R.N., and Yasar, D., 1999, Oscillating Quaternary water levels of the Marmara Sea and vigorous outflow into the Aegean Sea from the Marmara Sea–Black Sea drainage corridor: Marine Geology, v. 153, no. 1–4, p. 275–302, doi: 10.1016/S0025-3227(98)00078-4. Aksu, A.E., Hiscott, R.N., Mudie, P.J., Rochon, A., Kaminski, M.A., Abrajano, T., and Yasar, D., 2002a, Persistent Holocene outflow from the Black Sea to the eastern Mediterranean contradicts Noah’s Flood hypothesis: GSA Today, v. 12, no. 5, p. 4–10, doi: 10.1130/1052-5173(2002)012<0004: PHOFTB>2.0.CO;2. Aksu, A.E., Hiscott, R.N., Yasar, D., Isler, F.I., and Marsh, S., 2002b, Seismic stratigraphy of late Quaternary deposits from the southwestern Black Sea shelf: Evidence for non-catastrophic variations in sea-level during the last ~10000 yr: Marine Geology, v. 190, no. 1–2, p. 61–94, doi: 10.1016/ S0025-3227(02)00343-2. Algan, O., Gokasan, E., Gazioglu, C., Yucel, Z.Y., Alpar, B., Guneysu, C., Kirci, E., Demirel, S., Sari, E., and Ongan, D., 2002, A high-resolution seismic study in Sakarya Delta and submarine canyon, southern Black Sea shelf: Continental Shelf Research, v. 22, no. 10, p. 1511–1527, doi: 10.1016/S0278-4343(02)00012-2. Algan, O., Ergin, M., Keskin, E., Gökasan, E., Alpar, B., Ongan, D., and KirciElmas, E., 2007, Sea-level changes during the late Pleistocene–Holocene on the southern shelves of the Black Sea, in Yanko-Hombach, V., Gilbert, A. S., Panin, N., and Dolukhanov, P. M., eds., The Black Sea Flood Question: Changes in Coastline, Climate, and Human Settlement: New-York (USA), Springer, p. 603–631. Arkhangelskiy, A.D., and Strakhov, N.M., 1938, Geological structure and history of the evolution of the Black Sea: Izvestiia Akademii Nauk SSSR, Seriia Khimicheskaia, v. 10, p. 3–104. Arkhipov, S.A., Ehlers, J., Johnson, R.G., and Wright, H.E.J., 1995, Glacial drainage towards the Mediterranean during the middle and late Pleistocene: Boreas, v. 24, no. 3, p. 196–206, doi: 10.1111/j.1502-3885.1995 .tb00773.x. Bahr, A., Lamy, F., Arz, H., Kuhlmann, H., and Wefer, G., 2005, Late Glacial to Holocene climate and sedimentation history in the NW Black Sea: Marine Geology, v. 214, no. 4, p. 309–322, doi: 10.1016/j.margeo.2004.11.013. Bahr, A., Arz, H.W., Lamy, F., and Wefer, G., 2006, Late Glacial to Holocene paleoenvironmental evolution of the Black Sea, reconstructed with stable oxygen isotope records obtained on ostracod shells: Earth and Planetary Science Letters, v. 241, no. 3–4, p. 863–875, doi: 10.1016/ j.epsl.2005.10.036. Ballard, R.D., Coleman, D.F., and Rosenberg, G., 2000, Further evidence of abrupt Holocene drowning of the Black Sea shelf: Marine Geology, v. 170, no. 3–4, p. 253–261, doi: 10.1016/S0025-3227(00)00108-0. Bard, E., Hamelin, B., Fairbanks, R.G., and Zinder, A., 1990, A calibration of the 14C timescale over the past 30,000 years using mass spectrometric U-Th ages from Barbados corals: Nature, v. 345, p. 405–410, doi: 10.1038/345405a0. Benan, C.A.O.A., and Kocurek, G., 2000, Catastrophic flooding of an aeolian dune field: Jurassic Entrada and Todilto Formations, Ghost Ranch, New Mexico, USA: Sedimentology, v. 47, no. 6, p. 1069–1080, doi: 10.1046/j.1365-3091.2000.00341.x. Çagatay, M.N., Gorur, N., Algan, O., Eastoe, C., Chepalyga, A.L., Ongan, D., Kuhn, T., and Kuscu, I., 2000, Late Glacial–Holocene palaeoceanography of the Sea of Marmara: Timing of connections with the Mediterranean and the Black Seas: Marine Geology, v. 167, p. 191–206, doi: 10.1016/ S0025-3227(00)00031-1.

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The Geological Society of America Special Paper 473 2011

Rapid Holocene sea-level and climate change in the Black Sea: An evaluation of the Balabanov sea-level curve Ronald E. Martin* Department of Geological Sciences, University of Delaware, 103 Penny Hall, Newark, Delaware 19716-2544, USA Valentina Yanko-Hombach* Avalon Institute of Applied Science, 976 Elgin Avenue, Winnipeg, Manitoba R3E 1B4, Canada, and Department of Physical and Marine Geology, Odessa National I.I. Mechnikov University, 2 Shampansky per., Odessa, 65058, Ukraine ABSTRACT The investigation of rapid sea-level and climate change is critical to understanding the geologic history of the Black Sea and its effect on ancient civilizations of the region and adjacent areas. The current consensus of western scientists is that only local sea-level curves may be constructed because of local-to-regional changes in sedimentation, tectonics, and other factors. Recently, however, I.P. Balabanov published a synoptic sea-level curve for the entire Black Sea that spans the Pleistocene-Holocene transition and the Holocene based on older radiocarbon dates. This curve has been heavily criticized and is viewed skeptically by western workers for the reasons already mentioned as well as the use of questionable methodologies. Here, we examine Balabanov’s curve in light of these criticisms by comparing his sea-level curve to other independently derived sea-level and environmental indices. We find that, despite its drawbacks, many of the fluctuations of the Balabanov curve coincide with repeated ocean-atmosphere reorganizations, which involve shifts from cool to warm phases and corresponding changes in the species composition of foraminiferal assemblage ecozones, precipitation, and runoff. We suggest that following the initial invasion of the Black Sea by marine Mediterranean waters during the Pleistocene-Holocene transition, climatic amelioration (warming) following each cool phase of an ocean-atmosphere reorganization resulted in shifting precipitation patterns that produced repeated, rapid freshwater discharges into the Black Sea from surrounding rivers. In this scenario, runoff following each reorganization temporarily altered the species composition of foraminiferal assemblages, as noted in earlier studies. Freshwater discharges during the Holocene were likely lower than those envisioned by Balabanov but may have affected sea level sufficiently to alter coastal geomorphology and coastal aquifers rapidly, while causing the translocation of settlements from areas where submarine archaeological sites are now situated. Sea-level and climate change during the Pleistocene-Holocene transition may have been similar to that of the Holocene, but greatly amplified. *Martin—[email protected], Yanko-Hombach—[email protected]. Martin, R.E., and Yanko-Hombach, V., 2011, Rapid Holocene sea-level and climate change in the Black Sea: An evaluation of the Balabanov sea-level curve, in Buynevich, I.V., Yanko-Hombach, V., Gilbert, A.S., and Martin, R.E., eds., Geology and Geoarchaeology of the Black Sea Region: Beyond the Flood Hypothesis: Geological Society of America Special Paper 473, p. 51–58, doi: 10.1130/2011.2473(04). For permission to copy, contact [email protected]. ©2011 The Geological Society of America. All rights reserved.

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INTRODUCTION Several late Pleistocene and Holocene sea-level curves have been published during the past few decades for different sites along the coast of the Black Sea and adjacent waters. Representative curves for the northern Black Sea are shown in Figure 1. These curves occasionally agree in the timing, if not the magnitude, of the inferred sea-level fluctuations, but there is also substantial disagreement at other times. The disagreement between the curves has been attributed to a number of factors, which have been confounded with each other. The width and gradient of the shelves of the Black Sea vary substantially, potentially affecting the record of sea-level change. Extensive sediment input by the Danube, Dniester, Bug, and Dnieper Rivers has built the northwestern continental shelf outward in contrast to the narrow shelves east of Crimea and along the Turkish coast, where river input is minimal (Ivanova et al., 2007); this has resulted in differential sedimentation, subsidence, compaction, and erosion (e.g., delta lobe-switching) at different locales (e.g., Panin, 2009). The Black Sea is also tectonically active along both its southern margin (Koral, 2007; Shuisky, 2007; Yılmaz, 2007) and northern rim, where neotectonism due to mud volcanoes occurs (Alekseev et al., 2007; Huseynov and Aliyeva, 2007). Landslides and storms might also temporarily displace water onshore, leaving behind a local sedimentary record, which, if preserved, could mimic longer-term sea-level fluctuations (Kearney, 2001; N. Esin, 2009, personal commun.). The 14C dating methodologies previously used by Soviet scientists are also outdated relative to more recent accel-

erator mass spectrometer (AMS) techniques and are viewed skeptically. The potential effects of the indiscriminant collection of shells with little or no consideration of their taphonomic condition (S. Kroonenberg, 2007, personal commun.), the dating of bulk peat and shell samples (D. Kelterbaum, 2008, 2009, personal commun.), and variable reservoir effects on 14C dates must also be considered (Ryan, 2007; Nicholas et al., 2008; A. Chivas, 2008, personal commun.). Other effects that may impact the construction of sea-level curves include possible lead-lag effects between the Black Sea and adjacent water bodies as a result of the Bosporus Strait and Dardanelles, and anthropogenic activity related to deforestation and perhaps agricultural practices, such as the introduction of goats, the indiscriminant grazing of which accelerated erosion and the infilling of estuaries along the Aegean coast of Turkey (Kraft et al., 2003, 2007; Brückner, 2005; Brückner et al., 2006, 2010). Not surprisingly then, the Black Sea sea-level curves have met with increasing skepticism from investigators, especially western scientists. Consequently, some workers have suggested that locality-specific sea-level curves within the Black and Aegean Seas may be the only realistic option for deciphering the history of sea-level change and its effects on human settlement in the region (e.g., Kraft et al., 2003, 2007; Brückner et al., 2010). The Black Sea nevertheless presents a distinct advantage for reconstructing rapid Holocene sea-level and climate change because, with the exception of its connection to the Mediterranean through the Bosporus Strait and Marmara Sea, the Black Sea has remained isolated from other marine waters during the Holocene. As a result of its relative isolation, the Black Sea

Figure 1. Sea-level curves from the tectonically active region of the Caucasus (eastern Black Sea; A, B, C) and the passive margin of the northwestern Black Sea (D—near Dniester River; E—off Crimea). Figure is modified from Pirazzoli (1991, Plate 27: Black Sea II). Note that no large rivers flow into the Black Sea in the Caucasus and Crimea (see text for further discussion). Original sources for curves are: A—Balabanov and Izmailov (1988); B—Ostrovsky et al. (1977); C—Tchepalyga (1984); D—Voskoboinikov et al. (1982); E— Nevessky (1970).

Rapid Holocene sea-level and climate change in the Black Sea potentially amplifies the relatively subtle global climate signals of the Holocene by integrating signals over a vast catchment area of ~2.4 million km2 (Georgievski and Stanev, 2006) that spans much of western and central Europe (Stanev and Peneva, 2002). Within this catchment basin, the modern Danube, Dniester, and Dnieper Rivers alone account for ~85% of the runoff into the modern Black Sea (Likhodedova and Konikov, 2007). Most recently, Balabanov (2007) published a synoptic sealevel curve based on available older (uncalibrated) radiocarbon dates (Fig. 2). Balabanov’s curve would therefore appear to suffer from the same defects noted here. According to Balabanov (2007, and also personal commun.), 14C dates were constrained by extensive field observations, core data, geophysical profiling, topographic surveys, and paleogeographic reconstructions. Based on Figure 2, it appears that Balabanov used 14C dates of coastal-lagoonal peats as a lower boundary on the curve; such peats (unless they are basal peats, e.g., located immediately above noncompressible sands and gravels) are subject to compaction and stratigraphic displacement, as suggested by the wide scatter of data points for these environments. Like earlier curves (Fig. 1), Balabanov’s curve also exhibits a number of fluctuations, the timing and magnitude of which have been questioned at recent International Geological Correlation Programme (IGCP) 521– International Union for Quaternary Research (INQUA) 501 meetings in Odessa (Ukraine), Gelendzhik (Russia), Bucharest (Romania), and Izmir (Turkey). So, should we dismiss the available 14C dates and Balabanov’s sea-level curve, which is based on them, out of hand? Several of the local curves for the Eastern Mediterranean shown by Brückner et al. (2010, their fig. 3) exhibit a regression in the vicinity of 5 k.y. B.P., about when Balabanov’s curve also indicates a regression. Because it is a synoptic curve, Balabanov’s curve may therefore reflect eustatic (or at least large-scale regional) sea-level and climate change, the timing and magnitude of which may be obscured by the geologic setting at some, but not all, sites. Overlooking the possibility of such changes may lead us to ignore natural environmental change that significantly impacted ancient humans, impacts that may have had serious consequences for the development of modern humans. RESULTS AND DISCUSSION Timing of Sea-Level Fluctuations Like earlier Black Sea level curves (Fig. 1), Balabanov’s curve indicates a first-order rise of sea level into the early Holocene, followed by a much slower rate of rise, in agreement with global sea-level curves (Fig. 2). Superimposed on the first-order changes, smaller-scale fluctuations are noted. Several of the fluctuations in the Balabanov curve are paralleled by the sea-level curve of Filipova-Marinova (2007, based on coastal settings of Bulgaria) and the sedimentological and porewater analyses of Konikov (2007) for the northwestern Black Sea shelf (Fig. 2). Like some earlier sea-level curves (Fig. 1),

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Filipova-Marinova’s (2007) curve exhibits excursions above modern sea level, which some workers find anathema despite scattered reports of a mid-Holocene sea-level rise (e.g., most recently by Yu et al. [2009] and Kench et al. [2009] for atolls in the tectonically stable South China Sea and Maldives, respectively). Differential sedimentation and stratigraphic discontinuities certainly cannot be discounted in evaluating these curves, but temporal discrepancies between them fall within the range noted by Ryan (2007), who related his data to the Greenland Ice Core Project (GRIP) chronology in order to control for temporal variation of the reservoir effect. Several of the regressions in the Balabanov curve correspond to foraminiferal assemblage ecozone boundaries recognized by Yanko (1990; see also Yanko-Hombach 2007; Table 1; Fig. 2). The correspondence is not perfect, however. In some cases, Balabanov’s curve indicates a minor regression when Yanko’s data do not (4.7 k.y. B.P.) or Balabanov indicates a regression when the 14 C data appear to indicate otherwise (2.6 k.y. B.P.). Nevertheless, Figure 2 suggests that there is a relationship between the two data sets. Yanko (1990) and Yanko-Hombach (2007) inferred that the assemblage ecozone boundaries correspond to regressive sea-level phases because numerous ecozonal boundaries are inferred to correspond to disconformities or to sharp changes in foraminiferal diversity and abundance (see core descriptions in Yanko, 1990; Yanko-Hombach, 2007). The assemblages (and presumably sea-level fluctuations) appear to correspond to repeated invasions of Mediterranean species (or possibly ecophenotypes) accompanying eustatic sea-level rise and generally increasing marine conditions, because each assemblage ecozone contains distinctive Mediterranean taxa not found in earlier assemblage zones (Yanko-Hombach, 2007). This means that the influx of marine waters into the Black Sea during the Pleistocene-Holocene transition was a fluctuating phenomenon, not catastrophic as previously concluded by Ryan and coworkers (Ryan et al., 1997, 2003; Ryan, 2007). A similar trend in Black Sea salinity has been noted for dinocyst assemblages (Hiscott et al., 2007a, 2007b; Mudie et al., 2007; Marret et al., 2009) and calcareous nannofossils (Giunta et al., 2007). Furthermore, several of the fluctuations in the Balabanov curve and the regressions of Yanko’s (1990) work correspond to, or overlap with, ocean-atmosphere reorganizations of the Holocene, including the widely recognized climatic event at 8.3– 8.2 k.y. B.P. (Martin et al., 2007; Table 1; Fig. 2). The primary exception to this pattern is the reorganization from 3.2 to 2.4 k.y. B.P., which appears to have begun before any regression is indicated in the Balabanov curve (Fig. 2). The extent of the reorganizations is such that they appear to correspond to sea-level fluctuations detected in estuarine settings on either side of the North Atlantic. The sea-level fluctuations have been inferred from foraminiferal (ecostratigraphic) assemblage zones and are of parasequence scale (approximately several meters or more thick) along Delaware Bay, USA (Leorri et al., 2006; Martin et al., 2007); South Carolina estuaries (Colquhoun and Brooks, 1986; Colquhoun et al., 1995); the Seine estuary of northwestern France (Sorrel

Sea LevelFilipova-Marinova Uncalibrated

Sea LevelBalabanov

Sea LevelFilipova-Marinova Calibrated

Ky

Drier

Wetter

Molluscs & wood from bars, spits, terraces Alluvial-lacustrine molluscs Marine molluscs Lagoonal peats

Precipitation

Water Content

Water Content (%)

Figure 2. Holocene sea-level and climate change of the Black Sea. Foraminiferal assemblage zone boundaries of Yanko (1990), which indicate regressions, are indicated at bottom (see Table 1). Major (large boldface) and minor (small italicized boldface) ocean-atmosphere reorganizations are at top (from Mayewski et al., 2004); boldface italicized intervals indicate smaller range of reorganizations indicated by Mayewski et al. (2004, green bars of their fig. 1). Sea-level curve is from Balabanov (2007, his fig. 3). Time scale at top (k.y.) from Balabanov (2007, uncalibrated). Sea level curves from Balabanov (2007, uncalibrated) and Filipova-Marinova (2007, uncalibrated and calibrated, her figure 2 and table 1). Water content from Konikov (2007, calibrated, his figures 11, 13) and precipitation curve based on pollen studies of the Ukraine (N. Gerasimenko, 2008, personal communication, uncalibrated). Both calibrated and uncalibrated curves are shown for comparison. Ne—Neoeuxinian. See Table 1 for other abbreviations.

Sea Level (m)

54 Martin and Yanko-Hombach

Rapid Holocene sea-level and climate change in the Black Sea TABLE 1. TIMING OF HOLOCENE OCEAN-ATMOSPHERE REORGANIZATIONS AND BLACK SEA SEA-LEVEL CHANGES RECOGNIZED BY YANKO (1990) AND YANKO-HOMBACH (2007)* Ocean-atmosphere reorganization (k.y. B.P.)

Transgressions and regressions* (k.y. B.P.)

0.6–0.15 1.4–1.1

Recent (0.6–present) Korsunian (1.6–1.2) Nimphaean (2.2–1.6) Phanagorian (2.8–2.4) Dzemetinian (5.8–2.8) Eggrisian (6.4–6.2) Kalamitian (7.0–6.4)

3.2–2.4 4.2–3.8 5.9–5.3

8.4–8.0

Pontian (7.2–7.0) Vityazevian (7.6–7.4) Kolkhidian (8.2–7.6) Bugazian (9.4–8.2)

*Regressions shown in boldface.

et al., 2009); and the Arno River of Tuscany, Italy (Amorosi et al., 2009). Each reorganization is associated with widespread cooling followed by climatic amelioration, as inferred from ice-rafting events, the advance and retreat of alpine glaciers, strengthening and weakening of the westerlies, and changes in precipitation and aridity in lower latitudes (Mayewski et al., 2004). The reorganizations may well respond to other largerscale forcings such as solar insolation (Mayewski et al., 2004; Kroonenberg et al., 2007). Interactions of these systems with atmospheric dipoles such as the North Atlantic Oscillation (Martin et al., 2007), North Sea–Caspian Pattern (Gündüz and Özsoy, 2005), or El Niño–Southern Oscillation (ENSO), which affects climate in the nearby Caspian Sea region (Arpe et al., 1999), may also occur, but they are poorly understood (Bridgman and Oliver, 2006). Such reorganizations involving the behavior of the North Atlantic Oscillation may account for the Little Ice Age in Europe (Hurrell, 1995; Keigwin, 1996), which corresponds to a minor ocean-atmosphere reorganization and sea-level fall on the Balabanov curve (Fig. 2). Magnitude of Sea-Level Fluctuations Not only the timing, but also the magnitude of the sea-level fluctuations inferred by Balabanov has been questioned. There is little or no evidence, according to western workers, for sea-level fluctuations of several meters or more during the Holocene, at least in the Black Sea and Mediterranean regions. The waxing and waning of alpine glaciers in Europe during the Holocene were probably insufficient to force sufficient runoff and Black Sea sea-level change during the Holocene (Georgievski and Stanev, 2006). Instead, precipitation belts likely shifted back and forth over Europe (Kwiecien et al., 2009). Konikov et al. (2006) and Konikov (2007) found statistically significant discharge and sedimentation rates for the Dnieper River during the Holocene; they used the resulting periodicities to date regressions at 0.59, 3.08–2.18, 4.45–4.1, and 6.35–5.74 k.y. B.P. and transgressions

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at 1.18–1.24 and 5.4–5.45 k.y. B.P., corresponding to a number of sea-level fluctuations of the Balabanov (2007) curve. Similarly, based on more than 150 radiocarbon dates and 500 dendrochronologically dated oak trees, Starkel (2003) determined fluctuations of channel-forming discharges and sediment loads for the past 15 k.y. as far north as the Vistula River (which drains into the Baltic Sea), among them: 8.5–7.7, 6.5–6.0, 5.5–4.9, 4.4–4.1, 3.5–3.0, 2.7–2.6 k.y. B.P., 900–1150 A.D., and after 1500 A.D., which also overlap or coincide with a number of sea-level fluctuations of the Balabanov curve. This situation resembles that of the southeastern U.S. coast, where presumed sea-level fluctuations based on the elevation of midden sites may have resulted in part from the settlement of such sites during periods of higher precipitation and flooding, and perhaps storms (Colquhoun and Brooks, 1986; Colquhoun et al., 1995; Kearney, 2001). In the case of the Black Sea, Esin (2008, personal commun.; see also Esin and Kukleva, 2008) pointed out, based on numerical models, that the maximum sealevel rise resulting from freshwater inputs to the Black Sea could be no more than ~1 m during the Holocene, otherwise, freshwater would simply flow out the Bosporus. Nevertheless, it is possible that both freshwater input to (and outflow from) the Black Sea and eventual marine inflow from the Mediterranean may have been sufficient to promote continuous outflow from the Black Sea (as suggested by Hiscott et al., 2007a, 2007b). Panin et al. (2007; see also Borisova et al., 2006) calculated (based on river paleochannel dimensions) that freshwater discharge to the Black Sea was up to three times that of modern discharge rates during the PleistoceneHolocene transition. Freshwater outflow through the Bosporus that resulted from glacial retreat and freshwater discharge to the Black Sea during the late Pleistocene–Holocene transition then became increasingly attenuated as it gave way to eustatic sealevel rise, marine inflow, and increasing salinity (Lane-Serff et al., 1997). Brackish surface salinities resulting from continued (but increasingly attenuated) freshwater input might have produced density contrasts between surface and subsurface water masses sufficient to propagate internal waves that affected sedimentation and erosion on the Black Sea shelves (like that suggested by Ryan, 2007). Even if the effect of freshwater inputs to the Black Sea caused only limited sea-level change, minor sea-level fluctuations resulting from freshwater runoff may have affected human settlements in other ways. As sea-level rise slowed during the Holocene, deltas or other coastal settings may have prograded into some coastal areas (Coleman and Ballard, 2007), allowing the establishment of settlements. Following human colonization of coastal areas, freshwater discharges resulting in only a 1 m rise in the level of the Black Sea may have affected sea level sufficiently to alter coastal geomorphology and coastal aquifers, causing the translocation of settlements (although perhaps not on the scale envisioned by some workers; Turney and Brown, 2007) and accounting for the presence of submarine coastal archaeological sites at certain localities today.

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CONCLUSIONS Despite its possible drawbacks, the synoptic curve of Balabanov (2007) appears to reflect relatively subtle eustatic sea-level change related to hemispheric—and perhaps even global— climatic phenomena. The Black Sea amplified rapid, but subtle Holocene climatic signals because the basin is almost completely isolated, although the magnitudes of the sea-level oscillations are likely less than those proposed by Balabanov. Although the signals are subtle and may therefore often be overprinted by sedimentation, tectonics, etc., the sea-level and climatic oscillations reflected by them may have had significant impacts on human civilization. Following the initial invasion of the Black Sea by marine waters during the early Holocene, sea level in the Black Sea fluctuated repeatedly in response to both eustatic sea-level signals emanating from the Atlantic through the Mediterranean, Marmara Sea, and Bosporus Strait, and freshwater input from surrounding rivers, reflecting rapid shifts in the precipitation/ evaporation balance due to shifting wind patterns associated with ocean-atmosphere reorganizations involving one or more atmospheric dipoles. The decreasing amplitude of sea-level fluctuations suggests that both marine and freshwater influence on the Black Sea level waned through time, as would be expected as the overall climatic amelioration of the Holocene continued toward the present. Based on the inferences for the Holocene, freshwater outflow through the Bosporus that resulted from glacial retreat and massive freshwater discharge to the Black Sea during the late Pleistocene–Holocene transition became increasingly attenuated as it gave way to eustatic sea-level rise, marine inflow, and increasing salinity. Despite increasing attenuation during the Holocene, freshwater discharges may have affected sea level sufficiently to alter coastal geomorphology and coastal aquifers, causing the translocation of settlements and accounting for the presence of submarine archaeological sites at certain localities today. Given the apparent continuity of processes during the Holocene and Pleistocene-Holocene transition, reconnection between the Black Sea and marine waters of the Mediterranean appears to have been more oscillatory than catastrophic. ACKNOWLEDGMENTS This paper is an outgrowth of one initially presented at the 2007 Geological Society of America annual meeting and later at the 2007 IGCP 521–INQUA 501 meetings in Gelendzhik, Russia, Bucharest, Romania, and Izmir, Turkey. The senior author thanks the organizers for their invitations to present at these meetings and for the critical but constructive comments of other investigators at these meetings, especially those of Igor Balabanov, Helmut Brückner, Allan Chivas, Nikolay Esin, Natalia Gerasimenko, Daniel Kelterbaum, Evgeny Konikov, and Salomon Kroonenberg. Appreciation is also extended to the reviewers for constructive prepublication comments about the present paper. This

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Geoecology (GeoEcoMar), and Varna, Bulgaria, Department of Natural History of the Regional Historical Museum, p. 125–126. Ostrovsky, A.B., Izmaylov, Y.A., Balabanov, I.P., Skiba, S.I., Skryabina, N.G., Arslanov, S.A., Gey, N.A., and Suprunova, N.I., 1977, Novie dannie o palegidrologicheskom rezhime Chernogo moria v verkhnem pleistotsene i golotsene [New data on the paleohydrological regime of the Black Sea in the Upper Pleistocene and Holocene], in Kaplin, P.A., and Shcherbakov, F.A., eds., Paleogeografiia i otlozheniia pleistotsena iuzhnykh morei SSSR [Pleistocene Paleogeography and Sediments of the Southern Seas of the USSR]: Moscow, Nauka, p. 131–140 (in Russian). Panin, N., 2009, Contributions to understanding the sediment sink processes within the Danube–Black Sea system, in Gilbert, A., and YankoHombach, V., eds., Extended Abstracts of the Fourth Plenary Meeting and Field Trip of IGCP 521–INQUA 501 Project “Black Sea–Mediterranean corridor during the last 30 ky: Sea-level change and human adaptation,” 22–31 August 2009: Izmir, Kadir Has University, Dokuz Eylül University, and Çanakkale, Turkey, Çanakkale Onsekiz Mart University, p. 146–147. Panin, A.V., Sidorchuk, A.Yu., and Borisova, O.K., 2007, River runoff into the Black and Caspian Seas in the end of OIS-2, in Yanko-Hombach, V., Buinevich, I., Dolukhanov, P., Gilbert, A., Martin, R., McGann, M., and Mudie, P., eds., Extended Abstracts of the Joint Plenary Meeting and Field Trip of IGCP 521 “Black Sea–Mediterranean corridor during the last 30 ky: Sea-level change and human adaptation,” and IGCP 481 “Dating Caspian Sea Level Change,” 8–17 September 2007: Gelendzhik (Russia)– Kerch (Ukraine), Southern Branch of the Institute of Oceanology, Russian Academy of Sciences and Demetra Beneficent Foundation, p. 127–129. Pirazzoli, P.A., 1991, World Atlas of Holocene Sea-Level Changes: Amsterdam, The Netherlands, Elsevier, 300 p. Ryan, W.B.F., 2007, Status of the Black Sea flood hypothesis, in YankoHombach, V., Gilbert, A.S., Panin, N., and Dolukhanov, P., eds., The Black Sea Flood Question: Changes in Coastline, Climate, and Human Settlement: Dordrecht, The Netherlands, Springer, p. 63–88. Ryan, W.B.F., Pitman, W.C., III, Major, C.O., Shimkus, K., Moskalenko, V., Jones, G.A., Dimitrov, P., Görür, N., Sakınç, M., and Yüce, H., 1997, An abrupt drowning of the Black Sea shelf: Marine Geology, v. 138, p. 119– 126, doi: 10.1016/S0025-3227(97)00007-8. Ryan, W.B.F., Major, C.O., Lericolais, G., and Goldstein, S.L., 2003, Catastrophic flooding of the Black Sea: Annual Review of Earth and Planetary Sciences, v. 31, p. 525–554, doi: 10.1146/annurev.earth .31.100901.141249. Shuisky, Y., 2007, Climate dynamics, sea-level change, and shoreline migration in the Ukrainian sector of the circum-Pontic region, in Yanko-Hombach, V., Gilbert, A.S., Panin, N., and Dolukhanov, P., eds., The Black Sea Flood Question: Changes in Coastline, Climate, and Human Settlement: Dordrecht, The Netherlands, Springer, p. 251–277.

Sorrel, P., Tessier, B., Demory, F., Delsinne, N., and Mouazé, D., 2009, Evidence for millennial-scale climatic events in the sedimentary infilling of a macrotidal estuarine system, the Seine estuary (NW France): Quaternary Science Reviews, v. 28, p. 499–516, doi: 10.1016/j.quascirev.2008.11.009. Stanev, E.V., and Peneva, E.L., 2002, Regional sea level response to global climatic change: Black Sea examples: Global and Planetary Change, v. 32, p. 33–47, doi: 10.1016/S0921-8181(01)00148-5. Starkel, L., 2003, Palaeohydrology of central Europe, in Gregory, K., and Benito, G., eds., Palaeohydrology: Understanding Global Change: London, Wiley, p. 77–104. Turney, C.S.M., and Brown, H., 2007, Catastrophic early Holocene sea level rise, human migration and the Neolithic transition in Europe: Quaternary Science Reviews, v. 26, p. 2036–2041, doi: 10.1016/ j.quascirev.2007.07.003. Voskoboinikov, V.M., Rotar’, M.F., and Konikov, E.G., 1982, Sviaz’ ritmichnosti stroeniia tolshch golotsenovykh otlozhenii Prichernomorskikh limanov s kolebatel’nym rezhimom urovnia Chernogo moria [Relationship between the rhythmic composition of thick Holocene layers of the Black Sea region lagoons and the oscillatory level regime of the Black Sea], in Kaplin, P.A., Klige, R.K., and Chepalyga, A.L., eds., Izmeneniia urovnia moria [Sea-Level Fluctuations]: Moscow, Moscow State University, p. 264–274 (in Russian). Yanko, V.V., 1990, Stratigraphy and paleogeography of marine Pleistocene and Holocene deposits of the southern seas of the USSR: Memorie della Societa Geologica Italiana, v. 44, p. 167–187. Yanko-Hombach, V., 2007, Controversy over Noah’s Flood in the Black Sea: Geological and foraminiferal evidence from the shelf, in Yanko-Hombach, V., Gilbert, A.S., Panin, N., and Dolukhanov, P., eds., The Black Sea Flood Question: Changes in Coastline, Climate, and Human Settlement: Dordrecht, The Netherlands, Springer, p. 149–203. Yılmaz, Y., 2007, Morphotectonic development of the southern Black Sea region and the Bosphorus channel, in Yanko-Hombach, V., Gilbert, A.S., Panin, N., and Dolukhanov, P., eds., The Black Sea Flood Question: Changes in Coastline, Climate, and Human Settlement: Dordrecht, The Netherlands, Springer, p. 537–569. Yu, K.-F., Zhao, J.-X., Done, T., and Chen, T.-G., 2009, Microatoll record for large century-scale sea-level fluctuations in the mid-Holocene: Quaternary Research, v. 71, p. 354–360, doi: 10.1016/j.yqres.2009.02.003.

MANUSCRIPT ACCEPTED BY THE SOCIETY 22 JUNE 2010

Printed in the USA

The Geological Society of America Special Paper 473 2011

Global climate change and sea-level fluctuations in the Black and Caspian Seas over the past 200 years Eugen Konikov* Olga Likhodedova* Scientific and Educational Center of Geoarchaeology, Marine and Environmental Geology (SECGMEG), Palaeontological Museum, Odessa I.I. Mechnikov National University, 2 Dvorianskaia Str., Odessa 65082, Ukraine

ABSTRACT The problem of formation of the Black and Caspian Sea sea-level regime, which is considered as a component of Earth climate, is addressed through a time series of parameters for the past 200 years. To analyze the variables in sea-level changes, we applied modern methods of statistical processing: correlation, spectral and singularity analyses, and the wavelet analysis, among others. Using this approach, we prove that changes in climatic and hydrological parameters at global and regional scales are directly or indirectly reflected in sea-level regime. Based on statistical output, we propose a scenario of climatic sea-level changes for the short- and long-term future (until the end of 2100).

The scale of anthropogenic forcing on global climate change for the past five to ten decades is one of the most hotly debated questions in environmental science and politics. The data reported in the literature differ slightly (Fig. 1); however, all the graphs show an appreciable degree of warming during the 1930–1940s and over the last two decades We examined the following problems in our research: (1) changes in three climatic parameters (temperature, precipitation, and atmospheric pressure) during the instrumental measurement period; (2) the degree of influence of these parameters on river discharge and sea-level change in the Black and Caspian Seas by cross-correlation tests; (3) the periodic structure in time series using trend analysis, and spectral and wavelet analysis; and (4) reasons given for sea-level change of the world ocean, the Caspian and Black Seas, and the scenario of change predicted for the future of the Black Sea.

INTRODUCTION Earth’s climate is an extremely complex nonlinear system with numerous feedbacks, the dynamics of which are not obvious. Many publications contain inconsistent opinions concerning the causes of global warming. Central to these discussions is the question of the anthropogenic factor as a principal cause of the greenhouse effect, which lies at the root of global warming. Some authors argue that anthropogenic influences are already the dominant warming factor (Meleshko, 2007). Contrary to this opinion, the influence of greenhouse gases on climate change is considered to be not yet proven by some scientists, despite the strong consensus of the International Panel on Climate Change (IPCC, 2007) and Russian scientists (Yegorov, 2007; Datsenko and Monin, 2004; Boichenko and Voloschuk, 2006).

*Konikov—[email protected]; Likhodedova—[email protected]. Konikov, E., and Likhodedova, O., 2011, Global climate change and sea-level fluctuations in the Black and Caspian Seas over the past 200 years, in Buynevich, I.V., Yanko-Hombach, V., Gilbert, A.S., and Martin, R.E., eds., Geology and Geoarchaeology of the Black Sea Region: Beyond the Flood Hypothesis: Geological Society of America Special Paper 473, p. 59–69, doi: 10.1130/2011.2473(05). For permission to copy, contact [email protected]. © 2011 The Geological Society of America. All rights reserved.

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Konikov and Likhodedova Figure 1. Published records of surface temperature change over large regions (IPCC, 2001). All time series were smoothed using a 13-point filter. The Brohan et al. (2006) time series are anomalies from the 1961–1990 mean (°C). Each of the other time series was originally presented as anomalies from the mean temperature of a specific and differing base period. To make them comparable, the other time series have been adjusted to have the mean of their past 30 years identical to that same period in the Brohan et al. (2006) anomaly time series. Data sources are as follows: Köppen (1881): tropics and temperate latitudes, land air temperature. Willett (1950): global land stations. Callendar (1961): 60°N to 60°S land stations. Mitchell (1963): global land stations. Budyko (1969): Northern Hemisphere land stations and ship reports. Jones et al. (1986a, 1986b) and Hansen and Lebedeff (1987): global land stations. Hansen and Lebedeff (1987): global land stations. Brohan et al. (2006): global land air temperature and sea-surface temperature data.

METHODS The statistical laboratory data processing included: trend analysis, time-series analysis, spectral and cross-correlation analysis (using the “Statistica” package), and wavelet-analysis (Box and Jenkins, 1970; Daubechies, 2003). Wavelet analysis allowed us to investigate a wide range of time-domain periodicities. The basis of this method was generated in the mid-1980s by Grossman and Morle (Daubechies, 2003) as an alternative to Fourier transform for the analysis of temporal/spatial domains. Wavelet analysis divides the analyzed climatic and hydrological parameters into their constituent waves and components of various scales, and also provides time-specific process information. Mathematical modeling and the forecast of time series were executed using the exponential smoothing method in the Statistica package. RESULTS Global and Regional Changes in Climate The main feature of climatic change in the twentieth century was the increase in global temperature. Figure 1 shows the rapid rise of global temperature from 1920 to 1940, some decrease in temperature from 1960 to 1970, and a further increase at the end of the century. In the hundred years from 1906 to 2005, the average global temperature rose by 0.74 ± 0.18 °C, with most of this rise taking place over the last two decades. The average rate of warming calculated for the past 50 years is 0.13 ±

0.03 °C per decade, which is 2.5 times faster than that for the last 100 years (Meleshko, 2007). For the 16-year period from 1990 to 2006, the average temperature on Earth increased by 0.33 °C (Rahmstorf et al., 2007). Boichenko and Voloschuk (2006) showed that there is an alternation of long periods of warming and cold snaps in the extratropical parts of the Northern Hemisphere during the last 100 years that can be considered as a component of natural variations. These authors analyzed average annual temperatures of the Northern Hemisphere from 1856 to 2002 and determined a warm-cold oscillatory period of 66 years, using spectral analysis. The ascending branch of the last 66 summer periods coincides with the end of the twentieth century and the first two decades of the twenty-first century. On the basis of this study, Boichenko and Voloschuk (2006) drew their conclusion about the prevalence of natural factors driving climate change. These results correspond to our findings from the northwestern Black Sea region. Over the past 122 years, the climate on the coast of the northwestern region of the Black Sea has shown an increased trend in mean annual surface temperature, but no significant change was observed in precipitation. The values at Odessa are summarized in Table 1, based on the trend-analysis module of the Statistica package. This analysis allows the processing of numbers of observations that are not stationary. The regional temperature changes in the subtropicaltemperate northwestern Black Sea are less pronounced compared to the global changes. This probably reflects the fact that global warming is amplified at high latitudes because of the importance of reduced albedo following removal of snow and sea ice. At

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TABLE 1. RESULTS OF THE TREND ANALYSIS (CONSIDERED COEFFICIENT) Size of angular factor Parameters Observed period with the amendment Global temperature, T (°C) 1854–1989 (N = 136) 0.76 ± 0.06 Temperature (Northern Hemisphere), T (°C) 1854–1989 (N = 136) 0.68 ± 0.06 Temperature (Odessa), T (°C) 1876–2004 (N = 122) 0.28 ± 0.09 Precipitation (Odessa) (mm/yr) 1867–2001 (N = 135) Not significant

midlatitudes, temperature changes are not as pronounced, particularly in coastal regions, where temperature is further modulated by proximity to deep-water marine basins that act as heat reservoirs. It is notable that from the middle of the 1920s, the global situation varied slightly: there is an insignificant increase in mean annual temperature and a substantial increase in precipitation. The corresponding change of climatic conditions in a northwest part of Black Sea (Odessa) was expressed as a significant trend in temperature, annual precipitation, air pressure at ground level (1.5–3.0 m), and relative sea level (RSL; Fig. 2) during this period. The same RSL changes were characteristic for changes in the level of the world ocean. From 1923 to 1958, the level of the world ocean rose 70 mm, with an average speed of rise equal to 2.3 mm/yr. The maximum increase of RSL was observed in the Arctic Ocean (2.6 mm/yr) compared to 1.9 mm/yr in the Atlantic Ocean, 0.9 mm/yr in the Pacific Ocean, and 0.6 mm/yr in the Indian Ocean (Kalinin et al., 1975). The estimation of angular factors in the first part of the considered period (till 20 years of the twentieth century) does not show significant trends. Values of factors for the second period have a significance value p < 0.05 as shown in Table 2, where it is seen that for the past 100 years, most of the considered factors show a positive trend (Table 2). The increases have been caused primarily by a positive trend in the discharge of freshwater into the sea and an increase in atmospheric precipitation over the seawater. These changes, in turn, may have been caused by changes in atmospheric circulation, particularly the strengthening of meridional circulation of the air masses (Girs, 1971; Datsenko and Monin, 2004). With the rise in global temperature, we observe that strengthening meridional circulation forms a positive trend in components of the water balance of the Black Sea. Here, sea-level change as a whole will be correlated with sea-level change in the world ocean (positive dependence; Fig. 2). The considered parameters produce a complex system of interconnected factors, which greatly complicates the analysis of relationships. Even to a greater degree than the temperature, the numerical values of atmospheric circulation over the Black Sea are connected with sea level, river discharge, and atmospheric precipitation. Water-Level Changes in the Black and Caspian Seas as Responses to Global Climate The world ocean is a system of an integrated kind reflecting changes of a global climate. Thus, changes in global tempera-

ture reveal changes in sea-level change of the world ocean. Figure 3 shows a comparison of time series for global temperature, world ocean level and Atlantic Ocean level at Brest, west coast of France (from Woodworth and Player, 2003). Interdependence between these factors is characterized by cross-correlation factors of 0.61–0.72 (Table 3). For the Black Sea, numerical values of atmospheric circulation, even to a greater degree than the temperature, are connected with considered parameters such as sea level, river discharge, and atmospheric precipitation (Fig. 4). A stronger link with these indicators (temperature, pressure, and RSL) is found in the Caspian Sea. Analysis of the components of water balance in the Caspian Sea has revealed that the basic contribution (up to 72% of dispersion) in variability of sea level is attributed to inflow of river water within the Volga River basin (Mikhailov and Povalishnikova, 1998; Arpe and Leroy, 2007). The reasons for the change in the Volga discharge include variability in atmospheric precipitation (largely during the winter) in the river basin. The precipitation regime, in turn, can be defined by atmospheric circulation. It has been shown that an increase in sediment discharge into the Volga basin is related to sublatitudinal atmospheric circulation, and a reduction to a submeridional type of circulation. Other studies, however, have related changes in the Caspian precipitation regime to pressure systems in the Pacific Ocean (Arpe and Leroy, 2007). The original source of the moisture in the Volga basin includes influences of the North Atlantic Ocean. It is there that greater evaporation from the sea surface leads to an increase in the amount of moisture transferred to the Eurasian continent and, consequently, to increased atmospheric precipitation in the Volga basin. Recent water-level fluctuations in the Caspian Sea level have been influenced mainly by anthropogenic factors (Fig. 5). For example, there was a reduction in discharge because of irrevocable losses to in-filling sediment fill following dam construction on the water basin, evaporation from the surface of artificial reservoirs, and water extraction for irrigation. It is believed that since the 1940s, irreversible water consumption has steadily increased, which has led to reduction of inflow of river water to the Caspian Sea and an additional decrease in its level compared with the natural trend. At the end of the 1980s, the difference between actual sea level and the restored (natural) one has reached almost 1.5 m (Malinin, 1994). Thus, total water consumption in the Caspian Sea for those years is estimated at 36– 45 km3/yr (the Volga accounting for nearly 26 km3/yr). If not for the withdrawal of river water, the rise in sea level would have begun not in the late 1970s but in the late 1950s (Mikhailov and Povalishnikova, 1998).

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A

B

C

Figure 2 (continued on following page). Histograms (left column) and time series (right column) are shown for five factors (from top to down): (A) temperature, T (°C), (B) the sum of annual precipitation deposited (mm/yr), (C) atmospheric pressure at ground level (102 Pa), (D) Danube discharge (m3/s), and (E) sea-level marks (cm relative to Baltic level). Dark-blue color corresponds to data until 1920; red corresponds to data after 1920. Lines in the right part of drawing correspond to the theoretical normal law of distribution; lines in left correspond to the calculated trend.

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D

E

Figure 2 (continued).

This circumstance explains some of the discrepancy in the graphs presented in Figure 5. Results of our research (Konikov and Likhodedova, 2007; Likhodedova and Konikov, 2007, 2008) have shown the possible change in marine environmental processes occurring essentially simultaneously with changes in atmospheric circulation. This interdependence of level fluctuations in the Black and Caspian Seas is more strongly expressed in comparison with global temperature change, although undoubtedly, it is causally related to it. TABLE 2. RESULTS OF THE TREND ANALYSIS WITH SIGNIFICANCE VALUES P < 0.05 Parameters Significant value s of angular factor Temperature, T (°C) 0.24 ± 0.11 Precipitation (mm/yr) 0.34 ± 0.11 3 0.74 ± 0.08 Atmospheric pressure (10 Pa) 3 Danube discharge (m /s) 0.24 ± 0.09 Black Sea level (cm) 0.76 ± 0.08

Periodicity of Variations in Climatic Factors and Sea Level Wavelet analysis of a global temperature time series has allowed identification of cycles of around 10, 20–25, and 55– 60 years that are superimposed on longer periods (e.g., 100–130, 300–329-year cycles). These small-scale cycles are present throughout the entire period of recorded observations of 1864– 1984, including the industrial and postindustrial periods (Fig. 6). Most climatic indices show a dominant influence of the 60-year fluctuation (e.g., Datsenko and Monin, 2004). Boichenko and Voloschuk (2006) and Boichenko (2007) conducted spectral analyses of temperature for the interval from 1000 to 1850 A.D. and also found quasi-periodical fluctuations with periods of 57 ± 1 and 66 ± 2 years. The last observable 60-year cycle began in the 1970s. Its peak coincided with the beginning of the current twenty-first century. It is likely that the temperature maximum of the 60-year cycle has terminated, and some stabilization and the tendency for

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Figure 3. Correlation of global temperature anomalies (red) with sea-level change in the eastern Atlantic, at Brest, France (left) and the world ocean (right). Data are from Woodworth and Player (2003).

TABLE 3. INTERDEPENDENCE AMONG GLOBAL TEMPERATURE, WORLD OCEAN LEVEL, AND ATLANTIC OCEAN LEVEL (BREST, FRANCE) Parameter World ocean level Atlantic Ocean level Global temperature The maximum factor with The maximum factor with a a shift in 3 years shift in 1 year 0.72 ± 0.10 0.61 ± 0.09

Figure 4. Black Sea sea-level response (in cm) to changes in the direction of atmospheric circulation (A) and changes in sea level (B) at Odessa (dashed line) compared to the amplitude of the westerly circulation (W) with 11-year smoothing (solid line). E—easterly circulation; C—meridional circulation.

Global climate change and sea-level fluctuations in the Black and Caspian Seas

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Figure 5. Comparison of the Atlantic Ocean level at Brest to the level of Black Sea at Odessa (A) and between the world ocean and the level of the Caspian Sea (B). In the graph on the right, the y-axis is directed downward.

temperature decrease will be observed in the future. The assumption of a pause in global warming within the next decades, similar to the pattern from 1940 to 1970, has been forecast by Datsenko and Monin (2004). It is necessary to note that because the baseline precondition of anthropogenic scenarios consists of data observed in the second half of the twentieth century, the trend of influence of anthropogenic factors on a climate is accepted to be a unique variant that lasts into the twenty-first century. The basic deficiency of such approach was noted by Yegorov (2007), who showed that this model considers only a linear trend of natural changes and does not take into consideration that these changes can be component of more difficult, long ,and opposite-directed natural fluctuations,

as also pointed out by Ruddiman (2003). Agreeing with Yegorov (2007) as a whole, we do not reject completely the role of anthropogenic factors in climate change, but their role is obviously minor. The retrospective analysis of change of temperature for the past 5000 years (Fig. 7) allows comparison between modern warming and the small climatic optimum that took place during the Middle Ages–the Medieval Warm Interval (MWI) from ca. 1100 to 1350 A.D. A 3000-year record of Greenland ice sheet δ18O stable isotope data (IPCC, 2007) can be used to derive the changes in midannual temperature during this time. According to the spectral analysis of this time interval, there is evidence of distinct quasi-millennial, quasi–600-year and quasi–300-year cycles (Konikov, 2007a, 2007b; Shmuratko, 2001).

Figure 6. Cyclical changes in global temperature and wavelet transformation (left) and standardized wavelet values corresponding to scales 17 (~22-year period) and 29 (~60-year period).

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Figure 7. Variations of the isotope δ18O value from an ice core in Greenland (IPCC, 2007).

Possible Scenario of the Dynamics of the Black and Caspian Seas On the basis of the analysis of centuries-old processes of sea-level change in the Black and Caspian Seas, Shnitnikov (1969) concluded that their long-term fluctuations (their transgressions and regressions) should correspond to opposite phases of the general humidity. Periodicity of humidity is on the average equal to 1850 years. This basic pattern is shown in the Black Sea sea-level records and is also evident in shorter cycles of variability of humidity and water level in the seas. Comparison of the sea level during the time of instrument measurements shows this long-term cyclicity and also a quasi-centennial cycle (Fig. 8). Considering the 60-year fluctuations of temperature, in the next 20–30 years, some stabilization of water-level fluctuations of Black Sea is expected, with a slow trend toward its increase. For the Caspian Sea, after the increase period, after the middle of the 90-year cycle, a decrease is marked. The level of the Black Sea (on statistical forecasting) may rise 15 cm by 2050 A.D. and 30 cm by 2100 A.D. In the most adverse scenario, it could increase 40 cm by 2050 A.D. and 100 cm by 2100 (Konikov and Likhodedova, 2007). These estimations agree with the forecast by Pavlidis (2003) for the northern Atlantic and basin of the Arctic Ocean, where increase of RSL level will result in ~25 cm rise over 50 years. As a

Figure 8. Change of levels of the Black and Caspian Seas during the period of instrument measurement.

Global climate change and sea-level fluctuations in the Black and Caspian Seas whole, for longer periods, one possible scenario was provided by early computer-age forecast data of Shnitnikov (1969), as shown in Figure 9, based on palynology data and limnostratigraphy. At the boundary of the twentieth and twenty-first centuries, it is possible to judge global level inversion of the world ocean and Black Sea through the results of the periodicity derived from timeseries analysis of the Greenland δ18O isotope record for past 5000 years and from the reconstructed time series of Dnieper discharge for 4000 years (Fig. 10). From our data (Likhodedova and Konikov, 2006, 2007), it can be seen that the Dnieper River discharge is very strongly connected with Black Sea water-level fluctuations. The level of Caspian Sea is expected to rise in the long term, as predicted by 2000-year and 600-year periodicity of the climate. Superimposed on this background, short periodic changes occur in either phase or antiphase relation to the Black Sea. On the basis of the studied observations and interrelations between climatic factors and sea-level changes, statistical models and forecasts of development of the Black Sea water level have been constructed for the next decade. Mathematical modeling of the sea-level fluctuations is made under two schemes: (1) with a damped trend (a nonlinear fading trend), and (2) with a linear trend (Fig. 11). CONCLUSIONS The results of the research reported here reveal marked fluctuations of climatic parameters such as temperature, pres-

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sure, and atmospheric precipitation at global and regional levels, which are linked to features of atmospheric circulation in the Northern Hemisphere because of changes in prevailing directions of moving air and air-pressure systems (Sellers, 1969). The most notable changes in global temperature increase (global warming) and sea-level rise of the world ocean are primarily linked to periodic shifts of the natural processes, but are partially altered by anthropological and technological influences. By means of the statistical analysis, the relationship between relative sea level (RSL) in the world oceans and the midland Black and Caspian Seas is linked with climatic parameters. Consequently, we show the statistical dependence of Caspian Sea water level with RSL in the world ocean and changes in atmospheric circulation is greater than for Black Sea. Based on spectral and wavelet analyses for time series of climatic factors, river discharge, and sea levels for almost 200 years of measured intervals, statistically significant periods are determined to occur at cycles of 10, 20–25, and 55–60 years. Short periodic cyclicity has been established for intervals of the following lengths: seasonal, 2–3, and 5–7 years. Analysis of long data series for δ18O stable isotopes in Greenland ice, and Dnieper discharge reveal quasi–600-year and quasi–300-year cycles. We present the scenario of change in the climate and water level of the Black and Caspian Seas during the next few decades and the long-term forecast. Rise in the level of Black Sea water may proceed as early as 2020 to 2030 and will reach stabilization by the

Figure 9. Subcentennial variability of the total humidity and levels of the Baltic, Caspian, and Black Seas (from fig. 36 in Shnitnikov, 1969).

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Figure 10. Comparison of late Holocene changes in the Greenland δ18O isotope record and Dnieper discharge.

Figure 11. Comparison of sea-level values: measures of actual values are in blue, and the modeling results are in red. (A) Results using a nonlinear damped trend. (B) The model with a linear trend.

end of the twenty-first century. Then, a regressive decrease in level may be expected according to influence Shnitnikov humidity cycle. The Caspian Sea level will continue to rise, which corresponds to the model presented by Arpe and Leroy (2007).

and regional climate changes on formation of dangerous geological processes in the southwest of Ukraine.” The authors are grateful to all participants of the Project for assistance. REFERENCES CITED

ACKNOWLEDGMENTS This paper is made on the basis of the Project (Ministry of Education and Science) G/B-438 “Research of influence of global

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Köppen, W., 1881, Uber mehrjahrige Perioden der Witterung—III: Zeitschrift der Osterreichischen Gesellschaft für Meteorologie, v. XVI, p. 141–150. Likhodedova, O., and Konikov, E.G., 2006, Modeling of centennial Black Sea level changes as a basis for forecasting, in Yanko-Hombach, V., Buynevich, I.V., Gilbert, A., and Martin, A., Extended Abstracts of the Second Plenary Meeting and Field Trip of Project IGCP 521 “Black Sea– Mediterranean corridor during last 30 k.y.: Sea-level change and human adaptation,” 20–28 August 2006: Odessa, Ukraine, p. 111–114. Likhodedova, O., and Konikov, E.G., 2007, Analysis of sea-level changes in the Black Sea for the past 140 years and forecast for the future, in YankoHombach, V., Buynevich, I., Dolukhanov, P., Gilbert, A., Martin, R., McGann, M., and Mudie, P., eds., Extended Abstracts of the Joint Plenary Meeting and Field Trip of IGCP 521 “Black Sea–Mediterranean corridor during the last 30 k.y.: Sea-level change and human adaptation,” and IGCP 481 “Dating Caspian Sea level change,” 8–17 September 2007: Gelendzhik, Russia–Kerch, Ukraine, Southern Branch of the Institute of Oceanology, Russian Academy of Sciences and Demetra Beneficent Foundation, p. 106–109. Likhodedova, O., and Konikov, E.G., 2008, Sea-level fluctuations in the Black and Caspian Seas and global climate changes, in Yanko-Hombach, V., Buynevich, I., Dolukhanov, P., Gilbert, A., Martin, R., McGann, M., and Mudie, P., eds., Extended Abstracts of the Joint Plenary Meeting and Field Trip of IGCP 521 “Black Sea–Mediterranean corridor during the last 30 k.y.: Sea-level change and human adaptation,” and IGCP 481 “Dating Caspian Sea level change,” 8–17 September 2007: Gelendzhik, Russia– Kerch, Ukraine, Southern Branch of the Institute of Oceanology, Russian Academy of Sciences and Demetra Beneficent Foundation, p. 106–109. Malinin, V.N., 1994, Problema prognoza urovnya Kaspiyskogo moray [Problem of the Forecast of Caspian Sea Level]: RGGMI volume: 160 p. (in Russian). Meleshko, V.P., 2007, Poteplenie klimata: Prichini I posledstviya [Climate Warming: The Reasons and Consequences]: Chemistry and a Life, Volume 4 (in Russian). Mikhailov, V.N., and Povalishnikova, E.S., 1998, Yescho raz o prichinakh izmenenij urovnya Kaspiyskogo moray v XX veke [Once again about the reasons of Caspian Sea level changes in the XXth century]: Moscow, Vestnik MGU, ser. 5, Geography (Sheffield, England), v. 3, p. 35–38. Mitchell, J.M., Jr., 1963, On the world-wide pattern of secular temperature change, in Changes of Climate: Proceedings of the Rome Symposium Organized by UNESCO and the World Meteorological Organization, 1961: Paris, UNESCO, Arid Zone Research Series 20, p. 161–181. Pavlidis, J.A., 2003, Vozmozhnosti izmeneniya urovnya oceana v nachale tret’ego tisyacheletiya [Possible changes of level of ocean in the beginning of the third millennium]: Oceanology, v. 43, no. 3, p. 441–446 (in Russian). Rahmstorf, S., Cazenave, A., Church, J.A., Hansen, J.E., Keeling, R.F., Parker, D.E., and Somerville, R.C.J., 2007, Recent climate observations compared to projections: Science, v. 316, p. 709, doi: 10.1126/science.1136843. Ruddiman, W.F., 2003, The anthropogenic greenhouse era began thousands of years ago: Climatic Change, v. 61, no. 3, p. 261–293, doi: 10.1023/B:CLIM.0000004577.17928.fa. Sellers, W.D., 1969, A global climatic model based on energy balance of the Earth-atmosphere system: Journal of Applied Meteorology, v. 8, p. 392– 400, doi: 10.1175/1520-0450(1969)008<0392:AGCMBO>2.0.CO;2. Shmuratko, V.I., 2001, Gravitatsiono-rezonansnij ekzotektogenez [GravitationalResonance Exotectogenesis]: Odessa, Astroprint, 347 p. (in Russian). Shnitnikov, A.V., 1969, Vnutrivekovaya izmenchivost’ componentov obschej uvlazhnennosti [Intracentury Variability of Components of the General Humidity]: Leningrad, Nauka [The Science], 243 p. (in Russian). Willett, H.C., 1950, Temperature trends of the past century: Centenary Proceedings of the Royal Meteorological Society of London, p. 195–206. Woodworth, P.L., and Player, R., 2003, The permanent service for mean sea level: An update to the 21st century: Journal of Coastal Research, v. 19, p. 287–295. Yegorov, A.G., 2007, Solnechniy tsikl I dva regima mnogoletnego izmeneniya prizemnogo davlenia v visokikh i umerennikh shirotakh severnogo polushariya Zemli v zimniy period [Solar cycle and two modes of longterm change of ground pressure in high and moderate widths of Northern Hemisphere of the Earth during the winter period]: Reports of the Ukrainian Academy of Sciences, p. 402–407 (in Russian). MANUSCRIPT ACCEPTED BY THE SOCIETY 22 JUNE 2010 Printed in the USA

The Geological Society of America Special Paper 473 2011

Paleogeography of the Pontic Lowland and northwestern Black Sea shelf for the past 25 k.y.

Evgeny Larchenkov* Sergey Kadurin* Physical and Marine Geology, Odessa National University, 2 Shampansky Per, Odessa, 65058, Ukraine

ABSTRACT Analyses of marine sediment lithology, paleorelief, and depositional environments on the northwestern Black Sea shelf were used for paleogeographic reconstructions reflecting the time periods of 30–25 ka, 15.5–15 ka, 11–10 ka, 9 ka, and 4 ka. The landscape of 25 ka, when sea level was 87 m below present, consisted of three geomorphic elements: (1) a denudation plain incised by numerous rivers and uniformly dipping southward, (2) a late Pleistocene alluvial terrace plain, within which the valleys of the Dnieper, Dniester, and Danube formed a common alluvial plain, and (3) a low coastal delta plain. The subsequent sea-level rise of the Neoeuxinian sea-lake to –55 m (15.5–15 ka) and –37 m (11–10 ka) resulted in the flooding of deltaic lowlands and a large portion of the alluvial terrace plain containing the deeply embayed Dniester and Dnieper limans. After the Drevnechernomorian transgressive phase at 9 ka, the sea flooded almost the entire late Pleistocene alluvial terrace plain, and the Dniester and Dnieper limans were converted to open marine embayments. Through Kalamitian time (4 ka), the entire Chilia section of the Danube delta was flooded. A large tract of land existed in the region of the Tendra Spit and Odessa bank. Around 25 ka, forest landscapes were common for the highlands and valley slopes of rivers and gullies. A steppe zone occupied the alluvial plain, which dominated the landscape to the south. This study demonstrates that paleogeographic reconstructions may serve as a basis for (1) locating submerged ancient settlements and (2) constraining possible migration routes.

*Larchenkov—[email protected]; Kadurin—[email protected]. Larchenkov, E., and Kadurin, S., 2011, Paleogeography of the Pontic Lowland and northwestern Black Sea shelf for the past 25 k.y., in Buynevich, I.V., YankoHombach, V., Gilbert, A.S., and Martin, R.E., eds., Geology and Geoarchaeology of the Black Sea Region: Beyond the Flood Hypothesis: Geological Society of America Special Paper 473, p. 71–87, doi: 10.1130/2011.2473(06). For permission to copy, contact [email protected]. © 2011 The Geological Society of America. All rights reserved.

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INTRODUCTION

MATERIALS AND METHODS

The problem of the formation of the Black Sea over the last 25–30 k.y. has for many decades been the focus of interest for geologists and geographers, archaeologists, and experts in socioeconomics and political studies. Forecasting possible shoreline changes for the next 50–100 yr is an important task, and a successful solution is an essential prerequisite to long-term planning strategies for sea transport and recreational zones, as well as for construction in the shore belt. All of these can affect decisions in the further economic development of many countries in the Black Sea region. Political differences account for the fact that the northern and southern parts of the Black Sea were studied separately until nearly the end of the twentieth century. Only since the “Iron Curtain” fell have western scientists been given a chance to investigate the northern part of the Black Sea. Most of them, however, are unable to use the huge compilations of geological information on this area that have been accumulated over the past 100 yr by many prominent scientists who published in Russian. For this reason, the challenging hypothesis by Ryan and Pitman (1998) on Noah’s flood in the Black Sea was not checked against the majority of the well-known facts on the geological structure and history of the northwestern shelf (Yanko-Hombach et al., 2007). Critics of this hypothesis discovered a problem that had not been obvious before. Apparently, significant differences exist in the depth estimates of coeval shorelines for the last 25–30 k.y. in different parts of the sea. This circumstance still hampers the establishment of a unified scheme of late Pleistocene and Holocene shoreline dynamics for the entire basin. One possible solution to this problem might be obtained through a paleogeographic reconstruction based on multidisciplinary study of different features of the bottom relief and sediments, as well as of the faunal, floral, and pollen complexes from these sediments. Paleogeographic analysis permits a clear separation of continental and marine environments, and it provides an opportunity to reconstruct landscape dynamics during different phases of a sea’s transgression (Dolukhanov et al., 2009). It also enables an estimate of sea depth fluctuations within the coastal belt and provides a means of comparing possible scenarios of environmental transformation within the Black Sea region during this century. It is well known that the Neoeuxinian sea-lake appeared 25–30 k.y. ago and was the last isolated basin in the Black Sea area. In its northwestern corner, it was characterized by a very narrow shelf connected with the continental slope and rise. The basin’s former coastline has been traced by following thick, sandy, nearshore deposits that currently lie at a depth of –87 to –88 m. The subsequent sea-level rise produced a wide shelf in the northwestern Black Sea, and this led to a progressive shift from terrestrial and nearshore depositional environments to marine shelf settings. Seafloor landscapes of the northwestern Black Sea shelf have been formed on submerged pre-Holocene North Pontic alluvial plain paleorelief as a result of sedimentation and the development of benthic biocoenoses. This paper is aimed at a paleogeographic reconstruction of the area (Fig. 1) during the late Pleistocene and Holocene.

Paleogeographic reconstruction was based on detailed surveys of the northwestern Black Sea shelf, conducted at scales of 1:200,000 and 1:50,000 by Prichernomorgeologia State Regional Geological Survey (Avrametz et al., 2007; Podoplelov et al., 1973–1975; Sibirchenko et al., 1983). The shelf surface was subdivided according to morphological and statistical analyses of seafloor relief, lithological characteristics, and quantitative parameters of the Holocene sediments, and the locations of benthic biocoenoses. Based upon statistical treatment of (1) water depth, (2) thickness of Holocene sediments, and (3) percentage of silt and clay within the sediments from 1596 locations, various landscape areas were identified. Each landscape area has a similar distribution of statistical parameters. Using a geographic information system (GIS) framework, we were able to correlate the paleorelief surfaces with their corresponding shoreline positions, taking into account uncertainties resulting from the limited number of radiometric dates. In so doing, we assume that the main features observed in the modern seafloor relief of the northwestern Black Sea shelf correspond to the erosional denudation surface formed in the aftermath of the post-Karangat regression. Subsequent modifications have been essentially due to sediment accumulation and, to a lesser degree, erosion. Hence, the statistically confirmed steep slopes may be considered reliable indicators of former shoreline positions (Larchenkov and Kadurin, 2006). Analysis of stratigraphy, lithology, mineralogy, and sediment thickness from more than 400 vibracores along the northwestern shelf enabled us to identify the facies corresponding to various types of sedimentary environments. Based on these data, faunal complexes within the sediments and the geomorphology of the study area were used to reconstruct the marine and coastal environments, as well as general features of paleorelief for the intervals of 30–25 ka, 15.5–15 ka, 11–10 ka, 9 ka, and 4 ka. GEOLOGICAL SETTING In general, the shelf surface is a gently southward-tilted plain, but according to statistical analysis, there are several depth intervals where the surface is flat or relatively steep (Larchenkov and Kadurin, 2005b). The inner shelf plain is separated from the coast by an offshore erosional coastal slope, which extends down to 10–15 m with inclines up to 5°. Seafloor relief irregularity to 40 m depth is clearly marked by linear incised depressions and troughs. Depression bottoms are almost flat with gradients less than 1°; their slopes are 3°–4°, but in some places, they approach 7°. Orientations of most depressions are submeridional, but some sections are aligned latitudinally. The largest depression is the Dnieper trough, which extends like an elbow from the DnieperBug liman to 40 m depth. The Odessa sandbank, with a depth of 10 m, sits on the northern part of the depression (Fig. 2). A large West-Tendra rise separates the Dnieper trough from the Karkinit depression, where the depth is up to 30–35 m.

Paleogeography of the Pontic Lowland and northwestern Black Sea shelf Southwestward of the depression, other elements of the inner shelf bottom relief include: the Dniester rise, the Dniester trough, and the Budak rise, which is joined to the paleovalley and avantdelta of the Danube River. All rises and terraces at the 40–45 m depth interval are divided by a gentle (1°–2°) slope and extend as an arc from Zmeiny Island to Tarkhankut Cape to form a flat plain 20–80 km in width. The seafloor surface gradient increases to 2°–4° at depths greater than 60 m, and bottom relief irregularity is marked by a depth amplitude of 5–10 m. This area of the shelf possesses widths ranging from 6 to 36 km, and it stretches to the shelf edge at 100 m depth. However, the edge is deeper than 130 m near the Crimean Peninsula southwest of Kalamit Bay, and south of Zmeiny Island. Quaternary Deposits on the Pontic Lowland and Northwestern Black Sea Shelf These deposits have been studied by many researchers, and there are numerous publications, mainly in Russian (Shnyukov, 1985; Fedorov, 1978; Veklich and Sirenko, 1976; and many others). The Quaternary sediments of the Pontic Lowland include eolian, proluvial, diluvial, eluvial-diluvial, alluvial, lake-alluvial,

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deltaic, and liman sequences. They cover watersheds and slopes of valleys, and they build low terraces along rivers and estuaries, deltas, beaches, spits, and sand bars (Veklich and Sirenko, 1976). Loess is widespread also. The liman valleys are filled by Upper Quaternary and Holocene deposits, with thickness ranging up to 30–40 m. Black, dark-green, or gray-green clay, and clayey sand with shells of marine and freshwater molluscs are the most common sediments here. Besides these, there are gray compact clays with sand and shells, which are overlain by black and dark-gray colored mud (Gozhik et al., 1987). The composition of late Pleistocene deposits and their distribution on the northwestern Black Sea shelf reflect the change of marine environment and its disconnection with the ocean (Fig. 3A). Neoeuxinian alluvial, alluvial-deltaic, and liman deposits are everywhere on the recent shelf. Silt, sand, clay, and peat are common within the upper part of the sequences, but the bottom parts consist mostly of sand and gravel, with silt being relatively rare and pebbles very rare (Avrametz et al., 2007; Podoplelov et al., 1973–1975). Lacustrine Neoeuxinian sediments have not been found anywhere on the shelf above –36 m, but in deeper water, they cover all underlying deposits. The sediments consist

Figure 1. Map of the study area.

Figure 2. Landscape regions of the northwestern Black Sea shelf.

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Figure 3. Lithological maps of the (A) Neoeuxinian, (B) Drevnechernomorian (Old Black Sea), and (C) Novochernomorian (New Black Sea) sediments.

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mostly of sand, including shelly debris and silt; clay is very seldom seen on the shelf. Shell, shelly sand, and muddy shell are very common in Holocene sediments on the northwestern Black Sea shelf; however, mud dominates in sediments from depressions in the seafloor relief (Sibirchenko et al., 1983). Holocene deposits on the shelf are subdivided into Drevnechernomorian—or Lower Chernomorian (Old Black Sea)—and Novochernomorian—or Upper Chernomorian (New Black Sea)—horizons. Drevnechernomorian sediments can be separated into Bugazian and Vityazevian subhorizons. Bugazian formations on the main part of the shelf consist of shell and shelly debris with sand, silt, and clay layers (Fig. 3B). Sand deposits dominate near the early Holocene seashore, where their thickness can be up to 1.5–2 m. In general, sediment thickness varies from 0.2 to 1 m; however, it increases to 3–5 m in the relief depressions where mostly silt and clay accumulate (Kakaranza and Larchenkov, 2007). The sediments include freshwater or brackish-water molluscs dominated by Dreissena or Viviparus in some parts, but brackish-water Cardium can sometimes also be found. The typical ostracod assemblage consists of an admixture of freshwater forms (Avrametz et al., 2007). Vityazevian deposits, which conformably overlie the preHolocene and Bugazian ones, consist mostly of mud. Sand is confined to a narrow belt in submerged shoreline areas and on spits. The deposits vary in thickness between 0.2 and 1.0 m, reaching 3.5–6.5 m on trench slopes as well as at the bench bottom and the 10 m contour on the slope of Odessa bank. The mollusc and ostracod assemblages reflect a gradual transition from the brackish-water Caspian to the saline Mediterranean environment. Euryhaline species are common in these deposits, and semi-stenohaline species are sometimes encountered near the top. The ostracod assemblages also have a mixed character: fresh, brackish, and marine (Avrametz et al., 2007). Novochernomorian sediments on the northwestern Black Sea shelf usually are subdivided into Kalamitian and Dzhemetinian subhorizons. Kalamitian marine sediments consisting of oozes and shells are widespread on most parts of the shelf where they cover Lower Holocene deposits, and sometimes preHolocene and pre-Quaternary deposits. There are remnants of nearshore sand accumulation forms, which represent evidence of post-Kalamitian sea-level decrease and erosion. As a rule, a layer of shells is common near the top of Kalamitian oozes in relatively deep places, and it changes to a thin layer of weathered and ground Mytilus detritus in shallow marine areas. This is an obvious sign of a hiatus at the end of Kalamitian time on the inner shelf. The thickness of the sediments in most cases is 0.3–1 m thick and reaches 2–4 m in trenches; it also reaches 2–4 m thickness in depressions (Avrametz et al., 2007; Gozhik et al., 1987). Widely distributed marine Dzhemetinian deposits consist mostly of shells, and rarely sands and gravel, with oozes largely in the central parts of depressions. Sand is accumulated as a relatively narrow strip along the coastal line. Only the central part of Odessa bank is characterized by a large area of sand deposits. As

a rule, areas containing shell deposits are located between oozes and sand areas (Fig. 3C). Dzhemetinian sediments usually consist of one layer, but occasionally, they are two- or multilayered. They unconformably overlie eolian-diluvial, alluvial, and lacustrine-alluvial sequences, and soils of the Lower, Middle, and lower part of the Upper Holocene, and sometimes continental Upper Quaternary rocks. The average thickness of Dzhemetinian deposits is 2–3 m, reaching 4–10 m in the central parts of depressions. The mollusc and ostracod assemblages in Novochernomorian sediments are characterized by the dominance of moderately stenohaline Mediterranean species combined with euryhaline ones (Gozhik et al., 1987; Ross and Degens, 1974; Shcherbakov et al., 1978a, 1978b). Assessment of Holocene sediment thickness on the northwestern Black Sea shelf through standard statistical analyses and GIS-aided mapping has established the spatial distribution of positive and negative thickness anomalies (Fig. 4) (Kadurin et al., 2008). Based on the law of distribution, sediment thickness has a logarithmic character. The bimodal nature of the histogram points to a nonuniform selection. The main factors responsible for the distribution of deposit thickness on the northwestern Black Sea shelf include topography, sedimentary setting, and tectonic movements, with antecedent topography being undoubtedly the key factor. We distinguished several seafloor regions, including bathymetric rises, depressions within the inner shelf, and the level surface of the middle and outer shelf (Larchenkov and Kadurin, 2007a). The thickness values were determined for the Sarat, Dniester, and Tendra rises, the Odessa sandbank, depressions of the Sarat, Dniester, and Dnieper River paleovalleys, and the middle to outer shelf region. For each seafloor region, the following statistical parameters were determined: mean thickness, standard deviation, variation coefficients of asymmetry, excess, mode, and scattering. Comparisons based on the Fisher criterion (uniformity of scattering) show that sediment thickness values correspond to three general categories: (1) rises, (2) paleovalleys, and (3) middle to outer shelf. The normal thickness for each region corresponds to a mean value with a standard deviation, with deviations in thickness considered as positive or negative anomalies (Table 1). RESULTS Dynamics of the Black Sea Level and Shoreline Changes over the Last 25 k.y. During the Neoeuxinian (30–25 ka), the Black Sea was an isolated semibrackish deep-water lake-sea. It could be connected with the world ocean only through the Mediterranean Sea, the level of which was, at that time, lower than that of the Neoeuxinian lake. The absence of a direct connection with the ocean allowed river drainage to have a prevailing influence on the lake level in pre-Holocene times.

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Figure 4. The spatial distribution of Holocene sediment thickness on the northwestern Black Sea shelf.

It is difficult to determine the spatial positions of ancient coastlines in the Black Sea. The reasons include (1) uncertainty in the determination of sediment features that demonstrate former land-sea transitional zones within bottom sediments, and (2) the extremely imprecise, and sometimes inaccurate, estimation of sediment age. The spatial and temporal position of submerged coastlines varies significantly in different parts of the Black Sea (Fig. 5). Among those identified are: –155 m (7.5–6.8 ka) on the Turkish shelf near Sinop (Ballard et al., 2000); –90 m (11.8 ka) on the Turkish shelf in front of Sakarya (Algan, 2003); –95 m (10.2– 8.6 ka) on the Romanian shelf (Lericolais et al., 2004); –100 m (Last Glacial Maximum, MIS-2) on the Romanian shelf (Wing-

uth et al., 2000); –120 m (13.4–11 ka) on the Romanian shelf (Ryan, 2003); –140 m (14.7–10.0 ka) on the northwestern shelf (between Karkinitsky and Kalamitsky Bays; Ryan et al., 1997); –100 m (20–18 ka) on the Crimean shelf (Shcherbakov et al., 1978a, 1978b); –100 m (ca. 800 ka; Krystev et al., 1990); and –147 to –70 m on the Caucasian, Bulgarian, and Kerch shelves (e.g., Glebov et al., 1996). Even the reconstructed depths of paleo–sea levels in the northwestern part of the sea do not compare favorably with those of the Kerch Peninsula and Caucasus area (Balabanov, 2007; Fedorov, 1988; Konikov, 2007). This might be the result of active tectonic movements in the area near the Caucasus. It should be emphasized that any active tectonic events, at least

TABLE 1. VALUES OF MEAN, AND POSITIVE OR NEGATIVE ANOMALIES IN HOLOCENE SEDIMENT THICKNESS FOR DIFFERENT REGIONS OF THE NORTHWESTERN BLACK SEA SHELF Sediment thickness (m) Region Mean Negative anomaly Positive anomaly Bathymetric rise 0.4–3.1 <0.4 >3.1 Depression 3.2–7.9 <3.2 >7.9 Middle to outer shelf 0.3–1.8 <0.3 >1.8

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Figure 5. Position of fossil shorelines according to various writers.

since Pliocene time, are not fixed (Aksu et al., 2002; Larchenkov and Kadurin, 2005a; Morgunov et al., 1981; Shcherbakov et al., 1978a, 1978b). This is very important for paleogeographic reconstructions. Therefore, depth levels of ancient zones of landsea transition determined with accuracy according to geological data can be considered as undisturbed paleo–sea levels (Larchenkov and Kadurin, 2008; Yanko-Hombach, 2006, 2007). However, complexity lies in the fact that late Pleistocene and Holocene transgressions occurred in pulsing, transgressive-regressive

stages. The coastline in a transgressive phase could be formed at the hypsometric level sometimes exceeding by 10–15 m the sea level established during the regressive phase of the given stage (Ivanov and Kakaranza, 2006). Obviously, more favorable conditions for preservation of a coastline that can be identified using sediment features existed during a regressive phase. It is common to divide the Neoeuxinian and Chernomorian (Black Sea) periods of transgression during the last 25–30 k.y. Each period has been divided into several transgressive-regressive stages (Table 2). The geological data obtained on the northwestern Black Sea shelf show evidence of a gradual rise in sea level, which is complicated by visible oscillations. Unfortunately, it is currently possible to estimate only the general duration of each transgressive stage; the duration of transgressive-regressive phases and the duration of stable sea-level intervals are unknown. This limitation prevents correct estimation of the rate of sea-level rise during each transgressive stage over the last 25 k.y. Using geological research data from “Prichernomorgeologia,” we determined the paleo–sea level position for the Neoeuxinian and Chernomorian (Black Sea) periods of transgression and their stages (Table 3). The average rate of sealevel rise was ~3 mm/yr. The maximal rate was not greater than 10 mm/yr during the 14–8.5 ka time interval. Calculated values for the average rate of sea-level change should probably be considered as underestimates.

Chernomorian

TABLE 2. STAGES OF NEOEUXINIAN AND CHERNOMORIAN TRANSGRESSIVE INTERVALS Time Depth of sea Transgressive interval Transgressive-regressive stage (ka) (m) Dzhemetinian

0–4.1

–2 to +2

Kalamitian

4.2–7.1

–7.5

Vityazevian

7.1–8.9

10.5–12.5

Bugazian

8.9–10.0

–25 to –22.5

Late Neoeuxinian

10.0–18.0

–37 to –35

Early Neoeuxinian

18.0–30.0

–55 to –57

Late Chernomorian

Early Chernomorian

Neoeuxinian

TABLE 3. CALCULATION OF AVERAGE RATE OF SEA-LEVEL RISE Amplitude of the Average rate of Time Sea-level position Time interval sea-level rise sea-level rise (ka) (m below present) (yr) (mm) (mm/yr) 0 0 1000 –1 00 0 – 1.0 1 +1 1000 3000 3 .0 2 2 1500 –4 00 0 – 2.7 3.5 +2 600 20 00 3 . 33 4.1 (±0.1) 0 (±2) 2550 4750 1.86 6.65 (±4.5) 4.75 (±2.75) 2050 6750 3.3 8.7 (±0.2) 11.5 (±1) 1550 12,250 7.9 10.25 (±0.25) 23.75 (±1.25) 1250 12,250 9.8 11.5 (±0.5) 36 (±1) 3500 19,000 5.43 14 (±1) 55 (±2) 10,000 31,000 3.1 24 (±1) 86 (±1)

Paleogeography of the Pontic Lowland and northwestern Black Sea shelf Shelf Landscapes of the Northwestern Black Sea Inner, central, and outer shelf areas have been delineated along the northwestern part of the Black Sea. The inner shelf is characterized by very rough relief, and it is bordered by a terrace at ~45 m depth. The central shelf area is very flat and situated at up to 60 m depth; it is separated from the outer shelf by a discernible seafloor ledge. According to the distribution of statistical parameters describing water depth, Holocene sediment thickness, and percentage of silt and clay within the sediments, the following landscape areas have been delineated (Larchenkov and Kadurin, 2007a): erosional coastal offshore slope, Danube avant-delta and paleovalley, depressions of river paleovalleys, terraces of the inner shelf, central shelf plain, and the outer shelf plain (Fig. 2) (Table 4): 1. The erosional coastal offshore slope is located in two areas. It is a very narrow zone up to 10 m deep, covered by shelly detritus or muddy shelly detritus with highly productive Mytilus biocoenoses. 2. The Danube avant-delta and paleovalley are characterized by a dominance of clayey mud with low carbonate content; this landscape type ranges up to 30 m in thickness and contains polychaete biocoenoses. 3. Depressions of river paleovalleys occupy the Dnieper and paleo-Dniester troughs, as well as the Karkinit depression, where water depth increases from 23–30 m in paleovalley heads to 40–42 m in the seaward direction. Pelitic material within the sediments increases with sea depth, and muddy shell-rich sediments on the slopes of depression valleys change downward into shelly muds, and then into muds. Also, organic matter content in the sediments increases from 0.5%–0.9% to 1.8%–2.4%. Here, Mytilus is the dominant species, and polychaetes species are subdominant. 4. Terraces of the inner shelf include the Western Tendra, Dniester, and Budak rises, which are located in areas where average depths are between 17 and 23 m, but less than 30 m, except over the Odessa sandbank, which is a shallow region. Here, common sediments are shelly debris and detrital shelly sands. Active

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biogenic sedimentation can be observed on the crests of the terraces. These areas are dominated by Mytilus biocoenoses. 5. The central shelf plain is located at depths between 45 and 60 m. Mainly muddy shell-rich sediments with 1.6%–2.3% organic carbon are found here. Both Mytilus and Phaseolinus are the dominant species in the biocoenosis. 6. The outer shelf plain is located southward, at depths greater than 60 m and down to the shelf edge. Here, the biocoenosis visibly changes, and Phaseolinus is the dominant species, polychaetes species are subdominant, and Mytilus is absent. DISCUSSION Paleogeographic Reconstructions As of 30 ka, the northwestern shelf of the Black Sea was land, with the coeval coastline at a depth of 90 m; subsequently, the shelf was inundated as part of a closed, semibrackish, deepwater Neoeuxinian sea-lake possessing a very narrow shelf passing into the continental slope. The landscape consisted of three geomorphic elements (Larchenkov and Kadurin, 2007b): (1) a denudation plain incised by numerous rivers and uniformly dipping southward, (2) a late Pleistocene alluvial terrace plain, within which the valleys of the Dnieper, Dniester, and Danube Rivers formed a common alluvial plain, and (3) a lowland coastal delta plain (Fig. 6). By 30–25 ka, the erosional denudation plain encompassed the entire area of the modern land surface and the vast region to the south of the Tendra Spit and Odessa sandbank, which was defined by a wide, elbow-shaped Dnieper River valley. The plain occupies a Miocene–Pliocene peneplain formed by the deltaic plains of the paleo-Prut, the paleo-Dniester, and the paleo-Bug Rivers. Its surface dips to the south from 180 to 80 m and consists of loess deposits incised by numerous rivers. To the south, the late Pleistocene alluvial terrace valley occupied the main part of the modern shelf; it was characterized by erosional and depositional processes. Here, the Dnieper, Dniester, and Danube valleys widened substantially and formed a single alluvial plain separated by local drainage divides. The upper reaches of the valley were

TABLE. 4. AVERAGE GEOLOGICAL CHARACTERISTICS OF BOTTOM SEDIMENTS IN DIFFERENT GEOMORPHIC ELEMENTS Thickness of Geomorphic elements Depth Gravel Sand Silt Mud marine deposits of the relief (m) (%) (%) (%) (%) (m) Coastal offshore 7.24 7.24 18.40 22.33 28.82 30.12 West-Tendra rise 1.28 1.28 55.38 30.47 7.79 5.02 Odessa sandbank 2.75 2.75 25.24 33.86 21.12 18.68 Dniester rise 1.83 1.83 53.83 30.82 5.58 8.45 Budak rise 2.62 2.62 49.17 37.79 6.71 5.05 Karkinit depression 0.56 0.56 52.72 19.16 15.80 8.19 Dnieper trough 5.95 5.95 13.73 13.96 33.68 31.89 Dniester trough 6.11 6.11 73.42 18.08 2.38 4.58 Danibe avant-delta 0.95 0.95 36.49 39.80 10.96 11.98 Central shelf plain 0.73 0.73 48.40 18.66 17.72 14.63 Outer shelf plain 0.23 0.23 27.74 25.15 22.14 24.51

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dominated by marshy floodplains with oxbow lakes. A low delta plain occupied what is now the present-day outer shelf, in some areas close to the shelf edge. It occupied a wide region of the nearshore zone. Its surface represented an actively forming alluvial plain with braided rivers confined by channel-margin bars. Vast lowlands formed between the bars and contained numerous lakes and probably inaccessible swamps. Sandy ridges and spits separated this region from the sea. The depositional environments of neritic offshore, avant-deltas, and very narrow neritic outer shelf were replaced by the bathyal upper part of the continental slope (Larchenkov and Kakaranza, 2007).

According to the pollen analysis (Gerasimenko, 2004; Veklich, 1989), it is possible to consider that forest and steppe landscapes developed in the area of the denudation plain watersheds. Forests occupied the heights and slopes of river valleys. The steppe zone spread southward where river valleys were the main element of the landscape. More than 20 types of soil existed in the area. As a result of water-level rise in the Neoeuxinian sea-lake by 15.5– 15 ka, the shoreline attained the –55 m mark. The early Neoeuxinian coastal lowland was submerged, along with more than one third of the alluvial plain. This relatively wide and shallow subhorizontal shelf setting had depths of 10–15 m (Fig. 7).

Figure 6. Paleogeographic scheme of the Pontic Lowland and northwestern Black Sea shelf for the beginning of early Neoeuxinian time (30–25 ka).

Paleogeography of the Pontic Lowland and northwestern Black Sea shelf Due to the rise in erosion baseline, the denudation plain experienced widening of the valleys and fluvial deposition. The alluvial terrace plain was still occupied by the composite Dnieper-Dniester valley and was dominated by extensive marshy floodplains up to 40 m in elevation. There were also the large paleo-Sarat and paleo-Kogilnik River valleys. Active deposition characterized the river valleys. The typical absence of deltas at the mouths of most rivers suggests a rather active hydrodynamic regime in the basin. The exception is the Dnieper River, which formed a large delta at its mouth. A wide distribution of Stephanodiscus diatoms in outer shelf sediments indicates fresh to brackish conditions.

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At the end of the Neoeuxinian phase (11–10 ka), sea level reached –37 m, and large parts of the alluvial terrace valley had been flooded. Dry land was limited to less than a quarter of its former area, with elevations typically exceeding 20 m. The shoreline was highly irregular due to the incursion of marine waters into numerous river mouths. Large Dniester and Dnieper limans were separated by a small drainage divide and extended deep into the alluvial plain, essentially separating it into different parts (Fig. 8). Another zone of the segmented alluvial plain was the branching system of the paleo-Sarat and paleo-Kogilnik valleys, which formed a large delta at their confluence. This region may have contained several islands. The estuary that formed in

Figure 7. Paleogeographic scheme of the Pontic Lowland and northwestern Black Sea shelf for the beginning of late Neoeuxinian time (15.5–15 ka).

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the flooded portion of the Danube valley bordered the plain to the southwest. Northward within the denudation plain, where elevations reached 140–160 m, widening river valleys continued to accumulate alluvium. Coastal accumulation forms developed along the wave-dominated portion of the inner shelf, with erosion and transit of sediments dominating the deep-sea trough. Domination of Stephanodiscus diatoms and relatively rare benthic species within the Neoeuxinian sediments (Gozhik et al., 1987) confirm the existence of fresh to brackish conditions. The existing pollen evidence (Gerasimenko, 2004) suggests that the vegetation of the erosion-depositional plain

consisted of forest-steppe on the watershed, and mixed forests (consisting of pine, birch, elm, oak, maple, and hornbeam and hazel as understory) were restricted to the slopes and bottoms of river valleys. Pollen complexes from Neoeuxinian deposits specify a dry and relatively cool climate. Thus, the periods of moderately damp and warm climate are marked. It is possible to consider that forest and steppe landscapes should have developed during all of Neoeuxinian time; however, only a small part of the steppe zone that had earlier spread to the south existed at the end of Neoeuxinian time. The development of river valleys was the main factor in landscape formation.

Figure 8. Paleogeographic scheme of the Pontic Lowland and northwestern Black Sea shelf for the end of late Neoeuxinian time (11–10 ka).

Paleogeography of the Pontic Lowland and northwestern Black Sea shelf The reconnection with the Mediterranean Sea at 9 ka led to stratification of the marine waters due to differences in salinity (Fedorov, 1983; Ivanov and Kakaranza, 2006). As a result of the Drevnechernomorian transgression, nearly all the late Pleistocene alluvial plain had been flooded. The Dnieper liman was converted into a large bay, which connected freely with the sea and extended almost to its modern shoreline. The Dniester liman was transformed into a wide open bay, with the Dniester River forming a large delta. The paleo-Sarat and paleo-Kogilnik Rivers developed a more expansive deltaic plain. A wide region of the denudation plain, with elevations above 20 m, remained to the south of the latitudinally oriented Dnieper valley (Fig. 9).

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The shelf was characterized by active hydrodynamic conditions that precluded the accumulation of sediment, which was transported to the shelf edge and deep-sea trough. Sapropel mud and, less commonly, clay accumulated under neritic conditions on the outer shelf. The pollen complexes from Drevnechernomorian sediments show a sharp decrease in coniferous plants, an insignificant increase in deciduous trees and forbs, and a substantial increase in Chenopodiaceae, as compared with Neoeuxinian deposits (Isagulova and Gerasimova, 1981). This testifies to some warming of the climate, including an increase in its dryness. During the Kalamitian stage (until 4 ka), the coastline reached the –7 m mark, and the sea covered the entire present-day Chilia

Figure 9. Paleogeographic scheme of the Pontic Lowland and northwestern Black Sea shelf for the end of Bugazian time (9 ka).

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lobe of the Danube delta, with possible islands and promontories still existing in the area of the modern Dniester bank. A large exposed region around the Tendra Spit and present-day Odessa bank separated a wide, latitudinally oriented embayment of the lower Dnieper valley. Most of the familiar geomorphic elements had by then been established: the Balta sandy-muddy paleodelta denudation plain, the South Bug and Kogilnik erosion-denudation plains, the Budzhak erosional-depositional plain, and the Black Sea lowland. The Dniester-Tiligul loess plain at elevations from 150–160 m in the north to 45–50 m in the south was characterized by vast flat drainage divides, which correspond to segments

of the Pontic peneplain surface incises by deep (40–80 m) gullies and ravines. The undulating Danube-Dniester loess plain dips to the south from 15–16 m to 20–40 m and is segmented by a ravine-valley network from 80 m deep in the north to 30–40 m deep in the south (Fig. 10). The coastal plain of the Upper Pleistocene terraces extended in a 30–40-km-wide swath along the lower Danube and Black Sea coast. The initial level of its relief is the surface of the Upper Pleistocene alluvial, fluviolacustrine, and liman deposits of sand, gravelly sand, and sandy mud. The plain dips gently southward from 60–65 to 45–50 m. The terrace valleys of the Danube and

Figure 10. Paleogeographic scheme of the Pontic Lowland and northwestern Black Sea shelf for the end of Kalamitian time (4 ka).

Paleogeography of the Pontic Lowland and northwestern Black Sea shelf lower Dniester were produced by floodplain and terrace facies. Novochernomorian deposits cover the entire inner and middle shelf and are represented primarily by silt and clay oozes. Only within the Budzhak, Dniester, and Tendra highlands, as well as the Odessa bank, do sandy sediments dominate, and the thicknesses of deposits diminish to 3–5 times lower than in adjacent depressions. A vast field of shelly oozes has no apparent link with seafloor bathymetry. In the nearshore zone, wave action and longshore currents formed a variety of submerged accumulation forms, spits, and baymouth barriers. Analysis of Novochernomorian sediments shows the highest content of coniferous plant pollen during Holocene, some increase in Chenopodiaceae compared with Drevnechernomorian deposits, and an absence of pollen from deciduous plants and ferns (Cordova and Lehman, 2005; Gerasimenko, 1995). This might indicate a minor climatic cold snap during Novochernomorian time relative to the earlier Drevnechernomorian interval. The pollen spectrum from Holocene sediments of the inner shelf allows one to conclude that the adjoining land had been occupied by mixed forest, basically consisting of broad-leaved trees. Open areas were covered by mesophytic steppe coenoses, mainly consisting of forbs. At this time, the climate was relatively warm and humid. Possibly, all loess-covered watersheds supported subaerial steppe-forest and steppe landscapes; however, active sedimentation in river valleys produced landscapes on alluvial sand and mud.

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flooded by the sea, and the Dniester and Dnieper limans were converted to open marine embayments. During the Kalamitian stage (until 4 ka), sea level rose to –7 m, and the coastline was close to its current position. The sea covered the entire present-day Chilia lobe of the Danube delta, with possible islands and promontories still existing in the area of the modern Dniester bank. A large exposed region around the Tendra Spit and present-day Odessa bank separated a wide latitudinally oriented embayment of the lower Dnieper valley. Most of the well-known geomorphic elements of today were established. Later, after the Dzhemitinian stage, the line of the current seacoast was formed. Many limans were established, and all of them except the Dnieper-Bug became closed lakes. Seawater covered the Odessa bank. The coastal zone was segmented into accumulation and erosion areas, and the latter are the regions of active mass wasting. During the period from 25 ka until 11 ka, forest-steppe landscapes covered the drainage divides along the denudation plain, with forests common in the highlands and valley slopes of rivers and gullies. Southward, a steppe zone occupied the alluvial plain. The soil profile contained more than 20 soil types. Later, drainage divides consisting of loess still preserved subaerial steppe and forest-steppe landscapes, whereas riparian communities developed on alluvial sands and muds along river valleys undergoing active sedimentation. These paleogeographic reconstructions may serve as a basis for locating submerged ancient settlements dating to the past 7–8 k.y., as well as for constraining possible migration routes.

CONCLUSIONS ACKNOWLEDGMENTS Paleogeographic reconstructions for the time periods of 30–25 ka, 15.5–15 ka, 11–10 ka, 9 ka, and 4 ka have been based on analyses of continental and marine sediment lithology, depositional environments, and paleorelief on the northwestern Black Sea shelf. According to the distribution of statistical parameters of water depth, Holocene sediment thickness, and percentage of silt and clay within the sediments, the following landscape areas on the recent northwestern Black Sea shelf can be delineated: an erosional coastal offshore slope, a Danube avant-delta and paleovalley, depressions of river paleovalleys, terraces of the inner shelf, a central shelf plain, and an outer shelf plain. The landscape of 30–25 ka, when sea level was 87 m below present, was formed of three geomorphic elements that extended, in general, from northwest to southeast: (1) a denudation plain incised by numerous rivers; (2) a late Pleistocene alluvial terrace plain formed by the valleys of the Dnieper, Dniester, and Danube Rivers; and (3) a lowland coastal delta plain. During the transgressive Neoeuxinian period, a sea-level rise to –55 m (15.5–15 ka) and –37 m (11–10 ka) led to submergence of the entire deltaic lowlands and a large part of the alluvial terrace plain. The paleo-Dniester and paleo-Dnieper limans became deeply embayed, and the estuary that formed in the flooded Danube valley bordered the plain on the southwest. As a result of the Drevnechernomorian transgressive stage at 9 ka, almost the entire late Pleistocene alluvial terrace plain was

The paper has been based upon research conducted in connection with the International Geological Correlation Programme (IGCP) project 521 “Black Sea–Mediterranean corridor during last 30 ka: Sea level change and human adaptation.” Some of the data have been kindly provided by Prichernomorskoe State Regional Geological Survey—“Prichernomorgeologia.” A presentation based on these data was delivered during a session of the 2007 Geological Society of America (GSA) Annual Meeting, in Denver, Colorado, owing to the financial support from the GSA International Division. REFERENCES CITED Aksu, A.E., Hiscott, R.N., Yaşar, D., Işler, F.I., and Marsh, S., 2002, Seismic stratigraphy of late Quaternary deposits from the southwestern Black Sea shelf: Evidence for non-catastrophic variations in sea-level during the last ~10000 years: Marine Geology, v. 190, p. 61–94, doi: 10.1016/S0025-3227(02)00343-2. Algan, O., 2003, The connections between the Black Sea and Mediterranean during the last 30 k.y.: Geological Society of America Abstract with Programs, v. 35, no. 6, p. 461. Algan, O., Ergin, M., Keskin, S., Gökasan, E., Alpar, B., Ongan, D., and KirciElmas, E. 2007, Sea-level changes during the late Pleistocene–Holocene on the southern shelves of the Black Sea, in Yanko-Hombach, V., Gilbert, A.S., and Dolukhanov, P.M., eds., The Black Sea Flood Question: Changes in Coastline, Climate, and Human Settlement: Dordrecht, Springer, p. 603–631. Avrametz, V.M., Kakaranza, S.D., Sibirchenko, M.G., Mokrjak, I.M., Shvez, L.K., Makovetskaja, I.M., and Eremina, L.Iu., 2007, Zvit z provedennja

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geologichnoi ziomky masshtabu 1:200,000 pivnichno-zakhidnoi chastyny shelfu Chornogo morja v mezhakh arkushiv L-36-XIII. –XIV, XV. [Report on Geological Survey 1:200,000 within Card Segments L-36-XIII. L-36XIV, L-36-XV]: Odessa, Prychonomorske DRGP, 399 p. (in Ukrainian). Balabanov, I.P., 2007, Holocene sea-level changes of the Black Sea, in YankoHombach, V., Gilbert, A.S., Panin, N., and Dolukhanov, P.M., eds., The Black Sea Flood Question: Changes in Coastline, Climate, and Human Settlement: Dordrecht, Springer, p. 711–730. Ballard, R.D., Coleman, D.F., and Rosenberg, G.D., 2000, Further evidence of abrupt Holocene drowning of the Black Sea shelf: Marine Geology, v. 170, p. 253–261, doi: 10.1016/S0025-3227(00)00108-0. Cordova, C.E., and Lehman, P.H., 2005, Holocene environmental change in southwestern Crimea (Ukraine) in pollen and soil cords: The Holocene, v. 15, p. 263–277, doi: 10.1191/0959683605hl791rp. Dolukhanov, P.M., Kadurin, S.V., and Larchenkov, E.P., 2009, Dynamics of the coastal north Black Sea area in late Pleistocene and Holocene and early human dispersal: Quaternary International, v. 197, p. 27–34, doi: 10.1016/j.quaint.2008.04.002. Fedorov, P.V., 1978, Pleistotsen Ponto-Kaspiia [Pleistocene of the PontoCaspian]: Nauka, Moscow, Trudy Geologicheskogo Instituta AN SSSR 310, 166 p. (in Russian). Fedorov, P.V., 1983, Sredizemnomorskie transgressii v geologicheskoi istorii Chernogo moria [Mediterranean transgressions in the geological history of the Black Sea]: Bulleten Moskovskogo obshchestva ispytatelei prirody, Otdel geologicheskii [Bulletin of the Moscow Investigators of Nature, Geological section), v. 58, no. 6, p. 120–126 (in Russian). Fedorov, P.V., 1988, Problema izmeneniia urovnia Chernogo moria v pleistotsene [Issues of the Black Sea level changes in Pleistocene]: Bulleten Moskovskogo obshchestva ispytatelei prirody, Otdel geologicheskii [Bulletin of the Moscow Investigators of Nature, Geological section], v. 63, no. 1, p. 55–61 (in Russian). Gerasimenko, N.P., 1995, Holocene landscapes and climate changes in southwestern Ukraine, in Velichko, A.A., ed., Climate and Environment Changes of East Europe during Holocene and Late-Middle Pleistocene: Project Pole-Equator-Pole PEP 2: Moscow, Institute of Geography, Russian Academy of Sciences, p. 38–49. Gerasimenko, N.P., 2004, Rozvytok zonal’nykh ladshaftykh chetvertichnogo perioda na territorii Ukrainy [Development of Quaternary zonal landscapes in Ukraine]: Kiev, Naukova Dumka, 189 p. (in Ukrainian). Glebov, A.Y., Shimkus, K.M., Komarov, A.V., and Chalenko, V.A., 1996, Istoria i tendentsii evolutsii kavkazskogo regiona Chernogo moria [History and tendency of evolution of the Caucasian region of the Black Sea], in Glumov, I.F., and Kochetkov, M.V., eds., Antropogennoe zagryaznenie i protsess samoochischeniya kavkazskoy zony Chernogo morya [ManCaused Pollution and Process of Natural Self-Cleaning of Caucasus Zone of the Black Sea]: Moscow, Nedra, p. 28–56 (in Russian). Gozhik, P.F., Karpov, V.A., Ivanov, V.G., and Sibirchenko, M.G., 1987, Golotsen severo-zapadnoi chasti Chernogo moria [Holocene of the Northwestern Part of the Black Sea]: Kiev, Institute of Geological Sciences of the Academy of Sciences of the Ukrainian SSR, 46 p. (in Russian). Isagulova, E.Z., and Gerasimova, L.A., 1981, Sravnitelnaia palinologicheskaia kharakteristika poberezhia i dna severnoi chasti Chernogo moria [Comparison of palynological characteristics of onshore and bottom of the northern Black Sea], in Doklady IV Vsesojuznogo palinologicheskoi konferentsii [Reports of the IV Palynological Conference): Tjumen’, p. 61–62 (in Russian). Ivanov, V.G., and Kakaranza, S.D., 2006, Major stages of late Pleistocene– Holocene evolution of the northwestern Black Sea, in Yanko-Hombach, V., Buynevich, I., Chivas, A., Gilbert, A., Martin, R., and Mudie, P., eds., Extended Abstracts of the Second Plenary Meeting and Field Trip of Project IGCP 521 “Black Sea–Mediterranean corridor during last 30 k.y.: Sea-level change and human adaptation,” 20–28 August 2006: Odessa, Ukraine, Astroprint, p. 75–80. Kadurin, S., Eriomina, L.Yu., and Larchenkov, E., 2008, Distribution of Holocene sediment thickness on the northwestern Black Sea shelf, in Gilbert, A., and Yanko-Hombach, V., eds., Extended Abstracts of the Fourth Plenary Meeting and Field Trip of IGCP 521–INQUA 501 Project “Black Sea–Mediterranean corridor during the last 30 k.y.: Sea-level change and human adaptation,” 4–16 October 2008: Bucharest, Romania, National Institute of Marine Geology and Geoecology (GeoEcoMar), and Varna, Bulgaria, Department of Natural History of the Regional Historical Museum, p. 75–77.

Kakaranza, S.D., and Larchenkov, E., 2007, Litifatsii verkhnepleistotsen-golotsemovykh osadkov perekhodnoi zony ot severo-zapadnogo shelfa k glubokovodnoi vpadine Chernogo moria [Lithofacies of Upper Pleistocene and Holocene deposits in the transitional area from the northwestern shelf to the Black Sea deep-water depression]: Geologia i poleznye iskopaemye Mirovogo Okeana [Geology and Mineral Resources of World Ocean], v. 1, p. 89–99 (in Russian). Konikov, E.G., 2007, Sea-level fluctuations and coastline migration in the northwestern Black Sea area over the last 18 k.y. based on high-resolution lithological genetic analysis of sediment architecture, in Yanko-Hombach, V., Gilbert, A.S., Panin, N., and Dolukhanov, P.M., eds., The Black Sea Flood Question: Changes in Coastline, Climate and Human Settlement: Dordrecht, Springer, p. 405–436. Krystev, T.I., Limonov, A.V., Sorokin, V.M., and Starovoytov, A.V., 1990, Problema Chaudi bolgarskogo chernomorskogo shelfa [The problem of Chauda on the Bulgarian shelf of the Black Sea], in Krystev, T.I., ed., Geologicheskaia evoliutsia zapadnoi chasti Chernomorskoi kotlovini v neogen-chetvertichnoe vremia [Geological Evolution of the Western Part of the Black Sea in Neogene-Quaternary time]: Sofia, Institut Okeanologii Bolgarskoi Akademii Nauk, p. 349–361 (in Russian). Larchenkov, E., and Kadurin, S., 2005a, Vlijanie golotsenovykh tektonicheskikh dvizhenii na formirovanie donnykh landshaftov v zone Dunaj-Odessky zaliv [Influence of Holocene tectonic movements on sea bottom landscape formation in the Danube-Odessa Bay area]: Ekologia dovkillia i bezpeka zhittedijalnosti [Environment and Life Safety], v. 4, p. 13–19. Larchenkov, E., and Kadurin, S., 2005b, Modelling of coastline position along north-western part of the Black Sea for the past 25 k.y., in YankoHombach, V., Buynevich, I., Chivas, A., Gilbert, A., Martin, R., and Mudie, P., eds., Extended Abstracts of the First Plenary Meeting and Field Trip of Project IGCP 521 “Black Sea–Mediterranean Corridor during last 30 k.y.: Sea-level change and human adaptation,” 8–16 October 2005: Istanbul, Turkey, Kadir Has University, p. 102–103. Larchenkov, E., and Kadurin, S., 2006, Reconstruction of north-western Black Sea coastline positions for the past 25 k.y., in Yanko-Hombach, V., Buynevich, I., Chivas, A., Gilbert, A., Martin, R., and Mudie, P., eds., Extended Abstracts of the Second Plenary Meeting and Field Trip of Project IGCP 521 “Black Sea–Mediterranean corridor during last 30 k.y.: Sea-level change and human adaptation,” 20–28 August 2006: Odessa, Ukraine, Astroprint, p. 75–80., p. 105. Larchenkov, E., and Kadurin, S., 2007a, Northwestern Black Sea shelf bottom landscapes, in Yanko-Hombach, V., Buinevich, I., Dolukhanov, P., Gilbert, A., Martin, R., McGann, M., and Mudie, P., eds., Extended Abstracts of the Joint Plenary Meeting and Field Trip of IGCP 521 “Black Sea–Mediterranean corridor during the last 30 k.y.: Sea-level change and human adaptation,” and IGCP 481 “Dating Caspian Sea Level Change,” 8–17 September 2007: Gelendzhik (Russia)–Kerch (Ukraine), Southern Branch of the Institute of Oceanology, Russian Academy of Sciences and Demetra Beneficent Foundation, p. 102–104. Larchenkov, E., and Kadurin, S., 2007b, Paleogeographic reconstructions of Pontic Lowland and northwestern Black Sea shelf for the past 25 k.y.: Geological Society of America Abstracts with Programs, v. 39, no. 6, p. 429. Larchenkov, E., and Kadurin, S., 2008, Geological evidence for non-catastrophic sea-level rise in the northwestern Black Sea over the past 25 k.y.: Oslo, Norway, 33rd International Geological Congress, 6–14 August 2008. Larchenkov, E., and Kakaranza, S.D., 2007, Depositional environments on the northwestern Black Sea outer shelf and continental slope during late Pleistocene and Holocene, in Extended Abstracts of the Joint Plenary Meeting and Field Trip of IGCP 521 “Black Sea–Mediterranean corridor during the last 30 k.y.: Sea-level change and human adaptation,” and IGCP 481 “Dating Caspian Sea Level Change,” 8–17 September 2007: Gelendzhik (Russia)–Kerch (Ukraine), Southern Branch of the Institute of Oceanology, Russian Academy of Sciences and Demetra Beneficent Foundation, p. 99–101. Lericolais, G., Popescu, I., Panin, N., Guichard, F., and ASSEMBLAGE Scientific Team, 2004, Questions on the sea level fluctuations in the Black Sea since the Last Glacial Maximum: Assemblage Project, in YankoHombach, V., GÖRMÜŞ, M., McGann, M., Martin, R., John, J., and Ishman, S., eds., Abstract Volume of the Fourth International Congress on Environmental Micropalaeontology, Microbiology and Meiobenthology, 13–18 September 2004: Isparta, Turkey, p. 123. Lericolais, G., Popescu, I., Guichard, F., Popescu, S-M, and Manolakakis, L., 2007, Water-level fluctuations in the Black Sea since the Last Glacial

Paleogeography of the Pontic Lowland and northwestern Black Sea shelf Maximum, in Yanko-Hombach, V., Gilbert, A.S., and Dolukhanov, P.M., eds., The Black Sea Flood Question: Changes in Coastline, Climate, and Human Settlement: Dordrecht, Springer, p. 437–452. Morgunov, Iu.G., Kalinin, A.V., Kalinin, V.V., Kuprin, P.N., Limonov, A.F., Pivovarov, B.L., and Shcherbakov, F.A., 1981, Tektonika i istoria razvitia severo-zapadnogo shelfa Chernogo moria [Tectonics and Geological History of the Northwestern Black Sea shelf]: Moscow, Nauka Press, 244 p. (in Russian). Podoplelov, O.N., Karpov, V.A., Ivanov, V.G., Mokrjak, I.M., and Zagorodnij, G.G., 1973–1975, Otchet o geologicheskoi s’emki shelfa severo-zapadnoi chasti Chernogo moria masshtaba 1:50,000 [Report on Geological Survey 1:50,000 within the Northwestern Part of the Black Sea]: Odessa, Krymgeologiia, 367 p. (in Russian). Ross, D.A., and Degens, E.T., 1974, Recent sediments of Black Sea, in Degens, E.T., and Ross, D.A., eds., The Black Sea—Geology, Chemistry and Biology: American Association of Petroleum Geologists (AAPG) Memoir 20, p. 183–189. Ryan, W.B.F., 2003, New developments from continued explorations, in Program Abstracts of the International Conference “The Black Sea Flood: Archaeological and Geological Evidence,” 18–19 October 2003: New York, Columbia University, p. 5–6. Ryan, W.B.F., and Pitman, W.C., III, 1998, Noah’s Flood: The New Scientific Discoveries about the Event that Changed History: New York, Simon and Schuster, 320 p. Ryan, W.B.F., Pitman, W.C., III, Major, C.O., Shimkus, K., Moskalenko, V., Jones, G.A., Dimitrov, P., Görür, N., Sakınç, M., and Yüce, H., 1997, An abrupt drowning of the Black Sea shelf: Marine Geology, v. 138, p. 119– 126, doi: 10.1016/S0025-3227(97)00007-8. Shcherbakov, F.A., Koreneva, E.V., and Zabelina, E.K., 1978a, Stratigrafiia pozdnechetvertichnykh otlozhenii Chernogo moria [Stratigraphy of late Quaternary deposits of the Black Sea], in Pozdnechetvertichnaia istoria i sedimentogenez okrainnykh i vnutrennikh morei [Late Quaternary History and Sedimentogenesis of Outer and Inner Seas]: Moscow, Nauka Press, p. 46–51 (in Russian). Shcherbakov, F.A., Kuprin, P.N., Potapova, L.I., Polyakov, A.S., Zabelina, E.K., and Sorokin, V.M., 1978b, Sedimentatsiia na kontinental’nom shelfe Chernogo moria [Sedimentation on the Continental Shelf of the Black Sea]: Moscow, Nauka Press, 187 p. (in Russian). Shnyukov, E.F., 1985, Geologia shelfa Ukrainskoi SSR: Lithologia [Geology of the Ukrainian Shelf: Lithology]: Kiev, Naukova Dumka, 190 p. (in Russian).

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Sibirchenko, M.G., Karpov, V.A., Ivanov, V.G., Mokrjak, I.M., Zjultsle, V.G., and Tatarovskij, V.V., 1983, Otchet po izucheniiu litilogicheskogo sostava donnykh otlozhenii shelfa Chernogo moria s zeliju sostavleniia geologolitologicheskoi karty masshtaba 1:200,000 [Report on Lithological Study of the Black Sea Shelf for the Preparation of Geological-Lithological Map 1:200,000]: Odessa, Krymgeologia, 427 p. (in Russian). Veklich, M.F., and N.A. Sirenko, 1976, Pliotsen i pleistotsen levoberezh’ia Nizhnego Dnepra i ravninnogo Kryma [Pliocene and Pleistocene of the Left Side of the Lower Dnieper and Crimean Plain]: Naukova Dumka, Kiev (in Russian). Veklich, M.F., 1989, Etapy razvitia prirody Chernogo i Azovskogo morei v chetvertichnoe vremia i ikh koreliatsia s kontinentalnymi obstanovkami [Development stages of the Black and Azov Sea environments in the Quaternary and their correlation with continental environments], in Janshin, A.L., ed., Chetvertichny period. Paleogeografia i Litologia [Quaternary. Paleogeography and Lithology]: Shtiinza, Kishinev, p. 102–118 (in Russian). Winguth, C., Wong, H.K., Panin, N., Dinu, C., Georgescu, P., Ungureanu, G., Krugliakov, V.V., and Podshuveit, V., 2000, Upper Quaternary water level history and sedimentation in the northwestern Black Sea: Marine Geology, v. 167, p. 127–146, doi: 10.1016/S0025-3227(00)00024-4. Yanko-Hombach, V.V., 2006, Transformation of the Neoeuxinian lake into the Black Sea: Evidence from benthic foraminifera, in Yanko-Hombach, V., Buynevich, I., Chivas, A., Gilbert, A., Martin, R., and Mudie, P., eds., Extended Abstracts of the Second Plenary Meeting and Field Trip of Project IGCP 521 “Black Sea–Mediterranean corridor during last 30 k.y.: Sea-level change and human adaptation,” 20–28 August 2006: Odessa, Ukraine, Astroprint, p. 171–173. Yanko-Hombach, V.V., 2007, Controversy over Noah’s Flood in the Black Sea: Geological and foraminiferal evidence from the shelf, in Yanko-Hombach, V., Gilbert, A.S., Panin, N., and Dolukhanov, P.M., eds., The Black Sea Flood Question: Changes in Coastline, Climate, and Human Settlement: Dordrecht, Springer, p. 149–205. Yanko-Hombach, V., Gilbert, A.S., Panin, N., and Dolukhanov, P.M., eds., 2007, The Black Sea Flood Question: Changes in Coastline, Climate, and Human Settlement: Dordrecht, Springer, 999 p.

MANUSCRIPT ACCEPTED BY THE SOCIETY 22 JUNE 2010

Printed in the USA

The Geological Society of America Special Paper 473 2011

Nonpollen palynomorphs: Indicators of salinity and environmental change in the Caspian–Black Sea–Mediterranean corridor P.J. Mudie* Geological Survey Canada Atlantic, Box 1006, Dartmouth, Nova Scotia B2Y 4A2, Canada S.A.G. Leroy† Institute for the Environment, Brunel University, Uxbridge UB8 3PH, West London, UK F. Marret† Department of Geography, Liverpool University, Liverpool L69 7ZT, UK N.P. Gerasimenko† National Taras Shevchenko University of Kyiv, Earth Sciences and Geomorphology Department, Kyiv, GSP 680, Ukraine S.E.A. Kholeif† National Institute of Oceanography and Fisheries Egypt, Qayed Bay, Alexandria, Egypt T. Sapelko† Institute of Limnology, Russian Academy of Sciences, St. Petersburg 196105, Russian Federation M. Filipova-Marinova† Museum of Natural History, 41 Maria Louisa Blvd., 9000 Varna, Bulgaria

ABSTRACT Previous palynological studies of the Caspian–Black Sea–Mediterranean corridor primarily focused on pollen and spores for paleoecological and chronostratigraphic studies. Until recently, there has been less emphasis on the nonpollen palynomorphs, such as dinoflagellate cysts, algal and fungal spores, and animal remains. New studies of nonpollen palynomorphs in land-locked seas, estuaries, and lakes reported here indicate that they are important markers of salinity, nutrient loading, and human activity, including ballast discharge, farming, and soil erosion. We describe the nonoxidative laboratory processing methods necessary to extract nonpollen palynomorphs from marine- and brackish-water sediment samples. We list 48 nonpollen palynomorphs taxa from 37 surface sediments (including the past millennium) for cores

*Corresponding author, [email protected]. † Leroy—[email protected]; Marret—[email protected]; Gerasimenko—[email protected]; Kholeif—[email protected]; Sapelko— [email protected]; Filipova-Marinova—[email protected]. Mudie, P.J., Leroy, S.A.G., Marret, F., Gerasimenko, N.P., Kholeif, S.E.A., Sapelko, T., and Filipova-Marinova, M., 2011, Nonpollen palynomorphs: Indicators of salinity and environmental change in the Caspian–Black Sea–Mediterranean corridor, in Buynevich, I.V., Yanko-Hombach, V., Gilbert, A.S., and Martin, R.E., eds., Geology and Geoarchaeology of the Black Sea Region: Beyond the Flood Hypothesis: Geological Society of America Special Paper 473, p. 89–115, doi: 10.1130/2011.2473(07). For permission to copy, contact [email protected]. © 2011 The Geological Society of America. All rights reserved.

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Mudie et al. along the salinity gradient from <16‰ off the Danube Delta to >39‰ in the Aegean, Mediterranean, and Red Seas, for two Crimean saline lakes, the Caspian and Aral Seas, and for lakes in Iran and Kazakhstan. The main nonpollen palynomorphs taxa are illustrated and listed systematically to provide a baseline for future collaborative studies among Black Sea corridor palynologists. We outline the biological affinities of some nonpollen palynomorphs and discuss the initial results of the study in terms of what nonpollen palynomorphs may reveal about the history of the salinity in the Black Sea corridor and the impact of humans on soil erosion, plankton production, and harmful algal blooms.

INTRODUCTION Most previous palynological studies (e.g., Traverse, 1974, 1978; Rossignol-Strick, 1995; Aksu et al., 1995; Mudie et al., 2002; Filipova-Marinova and Christova, 2004) of the Black Sea and Eastern Mediterranean Sea focus on the study of pollen and spores for interpretation of climate change, because these terrestrial palynomorphs are continuously present throughout the Pliocene to Holocene sediments. However, there are many other kinds of palynomorphs in addition to pollen. These nonpollen palynomorphs (or NPM, nonpollen microfossils) include dinoflagellate cysts, various types of algal and fungal spores, zooplankton egg capsules, and other animal remains like the organic linings of small foraminifera (Van Geel, 2001). Nonpollen palynomorphs have often been discarded by Quaternary palynologists because of a lack of knowledge concerning their ecological and biological affinities; recently, however, new projects have focused on them because they contribute a wealth of information about paleoecological conditions, fires, and human land use (van Geel, 2006). Moreover, nonpollen palynomorphs are the focus of industry research on processes of black shale formation in anoxic basins, like the present Black Sea. Knowledge of nonpollen palynomorphs therefore represents a rich mine of information that should be better exploited, both for the completeness of Pleistocene and Holocene geological studies and for benefit of industry exploration. This research could represent a minimal effort to make the most use of the time already devoted to sampling, extracting, and counting the slides for pollen and spores. Furthermore, new studies of nonpollen palynomorphs in lakes and estuaries have shown that they are important markers of salinity, while others are more sensitive to nutrient loading and human activity (e.g., Brenner, 2001, 2006). In this paper, we document the most frequent nonpollen palynomorphs present in the surface samples from cores along the salinity gradient from <16‰ off the Danube Delta to >39‰ in the Aegean and Mediterranean Seas, in parts of the Caspian and Aral Seas, and in some saline or brackish-water coastal lakes adjoining the seas, and we also discuss what they might tell us about the history of the Caspian–Black Sea–Mediterranean corridor (abbreviated here as the Black Sea Corridor; Fig. 1). The main categories of nonpollen palynomorphs that have been reported in studies of Holocene sediments of the Black Sea corridor are dinoflagellate cysts (= dinocysts), including organic

linings of calcareous cysts, and various types of algae, including unicellular and colonial chlorophytes, cyanobacteria, filamentous zygnemataceans, and desmids. Also present are fungal spores and remains of fungal hyphae and fruiting bodies, zooplankton and insect skeletal remains, organic linings of small foraminifera and ostracods, scolecodonts (an old term for the mouthparts of marine worms), and other still unidentified forms known under the general name of incertae sedis or acritarchs. In some of the inland lakes and in the Nile Delta, filamentous cyanobacteria are found in the sediment deposits. In this paper, we present the initial results of studies made by members of the International Geological Correlation Programme (IGCP) 521 WG2 (Palynology), with the intent of stimulating further inputs from other palynologists working on nonpollen palynomorphs of the Caspian and Aral Seas, saline ponds (liman-lagoons), and other coastal environments of the Black Sea, Eastern Mediterranean, and the Nile Delta. At this early stage of our studies, we will simply demonstrate the existence of various nonpollen palynomorphs in the very large range of environments making up the Black Sea corridor, and discuss possible trends in relative abundances that seem to distinguish inland lakes with very unstable salinities from the continuously hypersaline Aegean and Levantine Seas. Key nonpollen palynomorphs taxa are illustrated, and where possible, we link their fossil names to modern life forms. At present, there is no standard application of terms describing salinity zones along the gradient from pure water (lacking any mineral salt) to brine ponds encrusted in sodium chloride (Fig. 2). For the purposes of this paper, therefore, we have adopted the terminology long used by scientists studying the benthic fauna of the Black Sea corridor (see Yanko-Hombach, 2007), but with a slight modification to accommodate the greater range of our salinity spectrum above 38‰. This terminology is also similar to that used in marine botany textbooks (e.g., Dawes, 1998). Figure 2A shows the salinity scale and the corresponding terminology for the planktonic nonpollen palynomorphs of the Black Sea corridor. The Yanko-Hombach scale for Black Sea corridor benthos is presented for comparison up to its maximum reported value (40‰). The accompanying map (Fig. 2B) shows the geographical distribution of the fresh to saline surface waters above the halocline at ~100 m depth. In the strongly stratified waters of the Black Sea corridor, the bottom water below ~200 m water depth is ~10‰ more saline than the surface layer. Bottom water in the brackish Black Sea is ~24‰ (brackish–semimarine), and it is in

Nonpollen palynomorphs the normal marine range (~29‰–34‰) below the semimarine (~21‰–24‰) surface water of the Marmara Sea. For comparison with our studies, the salinity ranges commonly used in coastal northwestern Europe and the Baltic Sea (Denys and de Wolf, 1999) are also presented. The salinity scale (Fig. 2A) shows that there is a much broader salinity range known as “brackish” in Europe (3‰–18‰–28‰) than in the Black Sea region (3‰–5‰–12‰), but there is close agreement in the definition of “freshwater” as less than 0.5‰. The misuse of the term “freshwater” for salinities less than 3‰–5‰ by sedimentologists and archaeologists (Turney and Brown, 2007) is unfortunate and confusing because permanent human habitation and irrigation in the Holocene Black Sea brackish lake would have been constrained by access to potable freshwater (<1‰), as discussed by Mudie et al. (2002, 2004). PREVIOUS STUDIES So far, most nonpollen palynomorphs studies have concentrated on peat (e.g., van Geel, 2001). From there, in the last decades, they have expanded to lake sediment. In Europe, a working group, ACCROTELM (Abrupt Climate Changes Recorded over the European Land Mass), has aimed to identify and publish

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on nonpollen palynomorphs of peat and lakes (www2.glos.ac.uk/ accrotelm/micproto.html). Much less work has been concerned with nonpollen palynomorphs of estuaries, closed seas, and open seas (Leroy, 1992; Matthiessen and Brenner, 1996; Brenner, 2001; Kholeif, 2004; Mudie et al., 2010). In the Black Sea corridor, several studies have been made of dinoflagellate cysts (dinocysts) and some of the acritarchs, starting with Wall et al. (1973) and Roman (1974). Traverse (1974, 1978) looked at dinocysts, acritarchs, fungal spores, and microforaminifera in Pliocene– Holocene Deep Sea Drilling Program (DSDP) cores. Mudie et al. (2001, 2002) studied the dinocysts, acritarchs, colonial algae, and fungal spores in the Black, Marmara, and Aegean Seas, while Filipova-Marinova and Christova (2004) and Atanassova (2005) reported on some Black Sea dinocysts, Cymatiosphaera and Pediastrum. Marinova and Atanassova (2006) described the nonpollen palynomorphs in a core from brackish (2‰–4‰) Lake Durankulak on the coast of northeastern Bulgaria. Hiscott et al. (2007) showed the Holocene history of microforaminiferal linings and Pediastrum in a core from the southwestern Black Sea shelf. Marret et al. (2004) and Leroy et al. (2007) reported on Caspian Sea dinocysts and other nonpollen palynomorphs, and Kazancı et al. (2004) listed various nonpollen palynomorphs from the coastal lagoon of Anzaleh (SW Caspian). Leroy et al. (2006) and

Figure 1. Regional map of the Caspian–Black Sea–Mediterranean corridor showing locations of the core sites used in this study. KBG— Kara-Bogaz Gol; Duran—Durankulak; Dz—Dzharylgach; Sa—Saki.

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Sorrel et al. (2006) also documented the dinocysts and other nonpollen palynomorphs from the Kara-Bogaz Gol and the Aral Sea, respectively, and Leroy et al. (2009) studied the nonpollen palynomorphs in Lake Sapanca east of the Marmara Sea. Giralt et al. (2004) reported on nonpollen palynomorphs in Lake Issyk-Kul, Kyrgyzstan. For the Aegean region, there is a report on the use of microforaminifera as an index of carbonate dissolution (Kotthof et al., 2008), and Aksu et al. (1995, 1999) and Zonneveld et al. (2001) studied dinocysts as indices of bottom-water oxidation

and as markers of plankton production and export to organic-rich sapropel sediments. The late Pleistocene–Holocene history of dinocysts in Nile cone sediments was documented by Kholeif and Mudie (2009), and initial studies of dinocysts and nonpollen palynomorphs on the inner Egyptian Shelf were reported by Marret et al. (2008). The main results from the pioneer studies of Wall et al. (1973) was a succession of dinoflagellate assemblages from one dominated by Spiniferites cruciformis and Pyxidinopsis psilata

Figure 2. Range of salinity in the Black Sea corridor. (A) Salinity scale showing the ranges of paleosalinity categories used for the Black Sea versus northwest Europe and the Baltic Sea (modified from Denys and de Wolf, 1999). (B) Map of the surface salinity in the study region; Aral Sea is presently hypersaline but was brackish for most of the past millennium.

Nonpollen palynomorphs (Tectatodinium psilatum), which they thought marked very low salinity, to a Lingulodinium assemblage, which they interpreted as indicating marine conditions at ca. 7 ka. This Lingulodinium cyst assemblage was followed by a peak of what they called an acritarch, Cymatiosphaera globosa, which they also attributed to marine influence. However, this interpretation was made just 5 yr after the link was established between living dinoflagellate plankton populations and fossil dinoflagellate cysts, and 2 yr after the biological link was made between Cymatiosphaera and the resting spore (phycoma) of the Chlorophyte alga Pterosperma (Martin, 1993). Since that time, laboratory cultures and more detailed ecological studies have shown that both these organisms, Lingulodinium and Pterosperma, respond to nutrient enrichment and water turbulence as much as to salinity. For example, Dale (1996) summarized the basic elements of the dinocyst salinity signal for the Baltic and Black Sea as consisting of: (1) coastal/neritic cyst species tolerating a broad range from normal marine (~35‰) to reduced marine conditions of ~20‰; (2) assemblages of Operculodinium centrocarpum and small Spiniferites spp. with reduced processes in salinities below 3‰; and (3) in the Black Sea, Pyxidinopsis psilata marking the interval 3‰–10‰, and above 7‰, an increase in abundance of Lingulodinium machaerophorum. Brenner (2001), however, found that in the entrance of the Baltic Sea and in the North Sea, there was a change from O. centrocarpum assemblages dominating below 15‰ to L. machaerophorum assemblages that may have been triggered by nutrient enrichment and by sensitivity to summer sea-surface temperature above 10 °C rather than just salinity. Similarly, Brenner (2001) found that the

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occurrence and abundance of P. psilata cysts in annual varves of Baltic Sea sediment is controlled by additional, as yet unknown, factors other than salinity. Consequently, simple conclusions about a marine index of salinity or a sea-level change from increases in Lingulodinium and Pterosperma, are not justified, and the trophic conditions of the waters must also be considered, in addition to stratification versus turbulence, nutritional requirements of associated heterotrophic species, and biological growth strategies (opportunistic r-type versus conservative K-type), as outlined for dinoflagellates by Taylor (1987). The terminology for trophic levels that we use in this paper is based on chlorophyll a (chl-a) values for the modern waters, as measured by satellite observations in June 1997 (Fig. 3) and the Aegean Sea trophic status classification of Ignatiades (2005). The succession of dominant nonpollen palynomorphs groups in the Black Sea cores studied by Wall and Dale (1974) looks more like a shift in phytoplankton from populations that characterize stenohaline oligotrophic conditions (e.g., Spiniferites cruciformis assemblages, according to Kouli et al., 2001) to a dominance of euryhaline opportunistic autotrophs such as L. machaerophorum, and followed by heterotrophs that characterize turbid nutrient- and diatom-rich water. Interpretation of the near-surface decline in Lingulodinium and Cymatiosphaera as a salinity decrease over the past 3 k.y. also disagrees with the occurrence of the exclusively marine coccolith Emiliania huxleyii in the Black Sea during this time (Marret et al., 2009). Studies of the carotenoid pigment isorerenieratene as a marker of stratification intensity, and δD records of C37

Figure 3. Distribution of trophic status categories for waters in the western study region, based on June 1997 values of chlorophyll a (in mg m3) and the classification of Ignatiades (2005).

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alkenones produced by haptophyte algae as a marker of surface salinity (van der Meer et al., 2008), however, suggest that a gradual increase in stratification and a 6‰ decline in salinity has occurred in the Black Sea over the past 2.7 k.y. This trend would be expected to favor the dominance of Lingulodinium, as evident from core-top assemblages off the Crimean Peninsula (Anna Gapanova, Odessa National University, 2009, personal commun.), where L. machaerophorum constitutes ~50%–90% of the core-top dinocyst assemblages. Alfred Traverse (1974, 1978) worked mainly on core cutter samples from DSDP cores back to the Miocene, and, therefore, his sampling intervals were more than a meter wide, providing no details for the late Pleistocene–Holocene. However, the long DSDP records do show the persistent presence of acritarchs, with fluctuations in their abundance that Traverse interpreted as indicating increased marine influence from sea-level changes. He devised a marine influence index (MI), calculated as MI = (dinocysts + acritarchs)/(dinocysts + acritarchs + pollen). Traverse was convinced that a high MI marked a sea-level transgression phase. It was a good idea, but the lack of correspondence with his marine dinocyst record reveals that it was not a very solidly based index because he did not know that some dinocysts and most acritarchs are fresh- and brackish-water species that respond to nutrient loading and temperature more than salinity or sea level. Beginning in 1992 (Leroy, 1992; Aksu et al., 1995), there has been a resurgence of interest in the study of nonpollen palynomorphs in the Black Sea corridor—mainly focused in dinocyst assemblages in the Marmara, Black, and Caspian Seas, as cited previously, but recently also including a wider range of nonpollen palynomorphs in the Nile Delta and cone, Red Sea coast, and Egyptian desert, and in the saline mountain lakes of northern Iran and Kyrgyzstan (see Fig. 2B; Table 1, and its references). LABORATORY PROCESSING METHODS AND SYSTEMATICS The palynological processing methods used by members of IGCP 521 for past studies include the acetolysis-KOH methods recommended by the ACCROTELM group, the cold HCl-HF digestion method preferred for marine sediments (Marret, 1993), and heavy liquid separation. Dale (1976) noted that hot HCl has an adverse effect on some dinocysts, and processing using only sonification and sieving is the best method when possible (although gypsum deposits or fluoride precipitate may necessitate HCl usage). In the past 30 yr, there have also been important advances in our knowledge about the preservation potential of brackish-water and marine nonpollen palynomorphs that must be considered when selecting a processing method (Versteegh and Blokker, 2004; Zonneveld et al., 2008). Although the walls of dinocysts and acritarchs are similar to pollen, they are not made of sporopollenin, which is very oxidation resistant. The walls of

organic dinocysts and chlorophyte phycomata are composed of more labile substances, produced by either the formation of linear carbon chains forming algaenans in the Chlorophytes, or by another pathway producing the aromatic wall chemistry of the Dinophyta. Versteegh and Blokker (2004) listed the biomolecular chemistry and resistance to acetolysis for a large number of algae that produce nonpollen palynomorphs. Postmortem polymerization of lipids is known to lead to fatty acid–based molecules in the walls of Eocene dinocysts, and there is still an important need for elucidating the chemical differences between the macromolecules produced by the living algae and their fossilized remains. To extract the full range of the nonpollen palynomorphs, strong oxidants like acetolysis and hot 10% KOH cannot be used. Figure 4 shows the recovery of dinocysts and selected algal nonpollen palynomorphs from five Black Sea cores when different processing methods are used. The low recovery of dinocysts in cores 72 and 1461 (probably also in Lake Durankulak, Table 1) reflects processing with the traditional KOH and acetolysis methods of Faegri and Iversen or Birks (Bryant and Wrenn, 1998) in contrast to the other samples that were processed by sieving, followed by chemical digestion with cold 10% HCL and warm 52% HF (e.g., Mudie et al., 2004). Other new methods use only sodium hexametaphosphate (NaPO3)6, as described by Riding and Kyffin-Hughes (2006), or dilute KOH (Bryant and Wrenn, 1998). Extraction using heavy liquid separation (e.g., Wood et al., 1996) is satisfactory in preventing oxidation damage of thin-walled palynomorphs, but it needs careful application to avoid loss of specimens by settling after coagulation of clumped grains. Hot HCl followed by cadmium-iodide heavy liquid separation was used to process the Crimean lake samples reported in this study. Zonneveld (2001) studied dinocysts in the western Levantine Sea and derived an index of oxidation sensitivity, showing that the most sensitive group is the protoperidinioid taxa. Cysts of heterotrophic, diatom-eating protoperidinioids are important markers of productivity in the Black Sea, but they are usually destroyed when acetolysis is used for processing (Marret, 1993). In contrast, L. machaerophorum is a moderately oxidationresistant taxon, so when acetolysis is used, inevitably this species appears in relatively high numbers, inflating its importance as an ecological marker. The Zonneveld oxidation index ranks organicwalled dinocysts as follows: (1) extremely sensitive: Protoperidinium spp., e.g., Brigantedinium, Quinquecuspis (formerly Multispinula), Echinidinium; (2) moderately sensitive: Operculodinium centrocarpum, Spiniferites species; (3) moderately resistant: Impagidinium, Nematosphaeropsis; and (4) resistant: Lingulodinium, Operculodinium israelianum, Polysphaeridium. The systematics used for the nonpollen palynomorphs are given in the following sections that describe the individual groups. Mainly, we followed the systematics used by Fensome et al. (1990) for acritarchs and prasinophytes, the dinocyst

Nonpollen palynomorphs

TABLE 1. OCCURRENCES OF NONPOLLEN PALYNOMORPHS (NPP) AND CORRESPONDING SURFACE-WATER SALINITIES (IN PARTS PER THOUSAND, ppt) OF LATE HOLOCENE SEDIMENTS (LAST 2–3 KA) FOR VARIOUS LACUSTRINE TO MARINE ENVIRONMENTS IN THE STUDY REGION EGYPT (1, 2) NILE (3) AEGEAN SEA (4) Core/Sample number: Qarun Del 6 Manz Hama Red S NC-1 NC2 22 G5 20 19 3 Surface salinity (ppt): 11 low 3 41.4 42 38 39 39.5 39.3 39.1 38.4 34 Pollen-Spores/g ×1000 80.1 70.5 + 20 A 8.69 24.3 0.11 0.06 0.21 1.23 0.37 Dinocysts/g ×1000 0 0 0 C C 7.232 6.221 0.12 0.13 0.37 0.16 0.37 Acritarch types: Leiosphere/Acritarch-8 Trav. 0 0 0 0 C + + 0 0 0 0 0 Chomotriletes minor C 0 0 0 R C C 0 0 0 0 0 Cymatiosphaera spp. C 0 0 C R O O 0 0 0 0 0 Hexasterias problematica 0 0 0 0 0 0 0 0 0 0 0 0 Micrhystridium O 0 0 R R + 0 0 0 0 0 0 Pseudoschizaea rubinus. 0 0 0 0 0 0 0 0 0 0 0 Pseudoschizaea sp. 0 0 + 0 + + 0 0 0 0 0 0 Pterospermopsis sp. 0 0 0 0 + 0 0 0 0 0 0 0 Sigmopollis spp. 0 0 0 0 0 0 0 0 0 0 0 0 Radiosperma corbiferum 0 0 0 0 0 0 0 0 0 0 0 0 O Tasmanites 0 0 0 C 0 0 0 0 0 0 0 Colonial algae 0 0 0 C 0 O 0 0 0 O 0 0 Botryococcus species 0 0 0 0 0 0 0 0 0 0 0 0 Pediastrum simplex + 0 + C R R 0 0 0 0 0 0 Pediastrum boryanum 0 0 C C 0 0 0 0 0 0 0 0 Cyanobacteria 0 0 0 0 0 0 0 0 0 0 0 0 Anaboena 0 0 0 0 0 0 0 0 0 0 0 0 Zygnemataceae 0 + + 0 0 R R 0 0 0 0 0 Fungal spores O O O 0 + 0 C 0 0 R 0 C Glomus-type A 0 0 C R C O 0 0 0 0 0 Neurospora 0 0 0 0 0 0 0 0 0 0 0 0 Sporormiella/Sordaria 0 0 0 0 0 0 0 0 0 0 0 0 Tetraploa 0 0 + 0 0 0 0 0 0 0 0 0 Thecaphora 0 0 0 0 0 0 0 0 0 0 0 0 Fungal hyphae C 0 0 C 0 C 0 0 0 0 0 0 O Fruiting bodies, germlings 0 0 0 0 0 0 0 0 0 0 0 Microforaminifera R 0 A C C C C C R C R C Tintinnids 0 0 0 0 0 + R 0 0 0 0 0 Thecamoebians 0 + + 0 + 0 0 0 0 0 0 0 Centropyxis species 0 0 0 0 0 0 0 0 0 0 0 0 cf. copepod eggs 0 + 0 0 C C A + 0 0 0 0 Rotifers 0 0 0 0 0 0 0 0 0 0 0 0 Scolecodonts 0 0 0 0 + 0 0 0 0 0 0 0 Artemia salina 0 O 0 0 C C C 0 0 0 0 0 Other animals 0 0 0 0 0 + + 0 0 0 0 0 Ostracod lining/jaw 0 0 0 0 0 R 0 0 0 0 0 0 Taxon diversity 10 5 8 12 17 19 12 4 3 5 3 4 Note: +—present; R—rare (<5%), O—occasional (6%–15%), C—common (16%–30%), A—abundant (>30%). Data sources for the sites, which are shown on Figure 1, are as follows. 1—Qarun Lake, Fayoum Depression: Kholief et al. (2007); Hamata mangrove swamp: Kholeif (2007); Suez tidal flat: Kholeif (2004). 2—Nile Delta site 6 and Manzala lagoon site 8: Leroy (1992). 3—Nile cone cores NC-1 and NC-2, Kholeif and Mudie (2009). 4—Aegean Sea sites 3–22 and G5, Marmara sites 2–12, and Black Sea sites 4–13: Mudie et al. (2002, 2004). 5—MAR02-45: Marret et al. (2009). 6—Sites 72* and 1461* of Roman (1974); site 1474 of Wall and Dale (1974). 7—Lake Dzarylgach: Sapelko and Subetto (2007). 8—Lake Saki: Gerasimenko and Subetto (2007). 9—Caspian Sea: Marret et al. (2004). 10—Leroy et al. (2007). 11—Lake Anzaleh (Anz.): Kazanci et al. (2004). 12—Kara-Bogaz Gol (KBG): Leroy et al. (2006). Iran: 13—Nowshahr: Ramezani et al. (2008). 14—Lake Urmia: Djamali et al. (2008); Aral*: Sorrel et al. (2006); Issyk-Kul (IssK*): Giralt et al. (2004). KYRG—Kyrgyzstan. (continued)

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TABLE 1. (continued) MARMARA SEA (4, 16, 17) 12 5 11 9 Sapa 23 22 22 20.1 0.5 3.2 2.73 5.78 3.89 A 2.53 6.09 7.87 2.33 R

BLACK SEA (4, 5 ,6, 7) Core/Sample number: 2 Dura 4 72* 1461* 45 13 7 1474+ Surface salinity (ppt): 24 3 17.6 15 17 18 22 22 22 Pollen-Spores/g ×1000 5.8 20 3.11 0.15 6 5.96 55.94 42 160 Dinocysts/g ×1000 2.7 R 2.97 0.5 0.12 3.99 2.07 12 2–62.0 Acritarch types: Leiosphere/Acritarch-8 Trav. 0 0 0 0 0 0 0 C + 0 R A A + Chomotriletes minor 0 0 0 0 0 0 0 0 0 0 0 0 0 0 O O Cymatiosphaera spp. 0 0 0 0 0 0 0 + R C A A Hexasterias problematica 0 0 0 0 0 0 0 0 0 0 0 0 0 0 Micrhystridium 0 0 0 0 0 0 C 0 0 0 0 0 0 0 Pseudoschizaea rubinus. 8 0 0 O R 0 0 O 0 0 A A A 0 O Pseudoschizaea sp. 0 0 0 0 0 + 0 0 0 0 0 0 0 Pterospermopsis sp. 0 0 0 0 0 0 0 0 + 0 + 0 0 0 O O Sigmopollis spp. 0 0 R 0 + C 0 0 C A A 0 Radiosperma corbiferum 0 0 0 0 0 0 0 0 0 0 0 0 0 0 Tasmanites 0 0 0 0 0 0 0 0 0 0 0 0 0 0 Colonial algae O 0 0 0 0 + A 0 R R R A A C Botryococcus species 0 0 0 0 0 0 R 0 0 0 0 0 0 0 Pediastrum simplex 0 0 0 0 0 + 0 0 0 0 0 0 0 0 Pediastrum boryanum 0 0 0 0 0 0 0 0 0 0 0 0 0 0 Cyanobacteria 0 0 0 0 0 0 0 0 0 0 0 0 0 0 Anabaena 0 0 0 0 0 + 0 0 0 0 0 0 0 0 Zygnemataceae 0 0 0 0 0 + R 0 0 0 0 0 0 0 Fungal spores R A A C C 0 C C + + C A A C Glomus-type O 0 0 0 0 0 + 0 0 0 + 0 0 0 Neurospora 0 0 0 0 0 0 R 0 0 0 0 0 0 0 Sporormiella/Sordaria R 0 0 0 0 + C 0 0 0 R 0 0 0 Tetraploa 0 0 0 0 0 + R 0 0 0 0 0 0 0 Thecaphora 0 0 0 0 0 + 0 0 0 0 0 0 0 0 Fungal hyphae 0 0 0 0 0 0 0 0 + 0 0 0 0 0 Fruiting bodies, germlings 0 0 0 0 0 0 0 0 + 0 0 0 0 0 Microforaminifera R A 0 A A + 0 C 0 0 A 0 0 0 Tintinnids 0 0 0 0 0 0 0 C 0 0 0 0 0 0 Thecamoebians 0 0 0 0 0 + 0 0 0 0 R 0 0 0 Centropyxis species 0 0 0 0 0 0 0 0 0 0 0 0 0 0 cf. copepod eggs O A 0 0 0 0 0 0 0 0 0 0 0 0 Rotifers 0 0 0 0 0 + 0 0 0 0 0 0 0 0 Scolecodonts R 0 0 0 0 0 0 0 0 0 0 0 0 0 Artemia salina 0 0 0 0 0 0 0 0 0 0 0 0 0 0 Other animals 0 0 0 0 0 0 0 0 0 0 R 0 0 0 Ostracod lining/jaw O 0 0 0 0 0 + 0 0 0 0 0 0 0 Taxon diversity 8 6 4 6 4 19 12 5 9 5 17 8 8 6 Note: +—present; R—rare (<5%), O—occasional (6%–15%), C—common (16%–30%), A—abundant (>30%). Data sources for the sites, which are shown on Figure 1, are as follows. 1—Qarun Lake, Fayoum Depression: Kholief et al. (2007); Hamata mangrove swamp: Kholeif (2007); Suez tidal flat: Kholeif (2004). 2—Nile Delta site 6 and Manzala lagoon site 8: Leroy (1992). 3—Nile cone cores NC-1 and NC-2, Kholeif and Mudie (2009). 4—Aegean Sea sites 3–22 and G5, Marmara sites 2–12, and Black Sea sites 4–13: Mudie et al. (2002, 2004). 5—MAR02-45: Marret et al. (2009). 6—Sites 72* and 1461* of Roman (1974); site 1474 of Wall and Dale (1974). 7—Lake Dzarylgach: Sapelko and Subetto (2007). 8—Lake Saki: Gerasimenko and Subetto (2007). 9—Caspian Sea: Marret et al. (2004). 10—Leroy et al. (2007). 11—Lake Anzaleh (Anz.): Kazanci et al. (2004). 12—Kara-Bogaz Gol (KBG): Leroy et al. (2006). Iran: 13—Nowshahr: Ramezani et al. (2008). 14—Lake Urmia: Djamali et al. (2008); Aral*: Sorrel et al. (2006); Issyk-Kul (IssK*): Giralt et al. (2004). KYRG—Kyrgyzstan. (continued)

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TABLE 1. (continued) CRIMEA (8, 9) Saki Dzar >40 100 A 2 0 0

CASPIAN (10, 11, 12, 13) ARAL IRAN (14, 15) KYRG Core/Sample number: C21 C14 Anz. KGB Aral* Nows Urm IssK* Surface salinity (ppt): 13 13 2–9 20–60 10 low 200 6 Pollen-Spores/g ×1000 15 10 50 2.–12 10.0–43 VA 5 34 Dinocysts/g ×1000 5 5 15 2.–18 2.0–10 0 + 0 Acritarch types: Leiosphere/Acritarch-8 Trav. 0 0 A 0 0 0 A 0 0 0 Chomotriletes minor 0 0 0 0 0 0 0 0 0 0 Cymatiosphaera spp. 0 0 0 0 0 0 A 0 0 0 Hexasterias problematica 0 0 0 0 0 0 R 0 0 0 Micrhystridium A + 0 0 0 0 C 0 0 0 Pseudoschizaea rubinus. 0 0 0 0 0 0 0 0 0 0 Pseudoschizaea sp. R 0 0 R 0 0 0 0 0 0 Pterospermopsis sp. 0 0 A O 0 0 0 0 0 0 Sigmopollis spp. 0 0 0 0 0 0 0 0 0 0 O O Radiosperma corbiferum 0 0 0 0 0 0 0 0 Tasmanites 0 0 0 0 0 0 0 0 0 0 Colonial algae 0 + A 0 0 0 R 0 0 + Botryococcus species 0 0 0 + R R-C A 0 0 A Pediastrum simplex 0 0 0 + R-O 0 + 0 0 0 Pediastrum boryanum 0 0 0 0 0 0 0 0 0 R Cyanobacteria 0 0 0 0 0 0 0 0 0 + Anabaena 0 0 A 0 O-A + 0 0 0 R Zygnemataceae O 0 0 0 + R-C 0 0 0 + Fungal spores O R 0 0 R 0 0 C 0 R-C Glomus-type 0 R 0 0 0 0 0 0 0 R Neurospora 0 0 0 0 0 0 0 0 0 0 Sporormiella/Sordaria 0 0 0 0 0 0 0 0 0 0 O Tetraploa 0 0 0 0 0 0 0 0 0 Thecaphora 0 0 0 0 0 0 0 0 0 R Fungal hyphae 0 0 0 0 0 0 0 0 0 0 Fruiting bodies, germlings 0 0 0 0 0 0 0 0 0 0 Microforaminifera 0 R 0 0 0 R 0 0 0 R Tintinnids O 0 0 0 0 0 0 0 0 0 Thecamoebians 0 0 0 0 0 0 0 R 0 + O Centropyxis species 0 0 0 0 0 0 0 0 0 cf. copepod eggs 0 0 0 0 0 0 0 0 0 0 Rotifers 0 0 0 0 R 0 0 0 0 0 Scolecodonts O 0 0 0 0 0 0 0 0 0 Artemia salina O 0 0 0 0 0 0 0 0 0 Other animals 0 + 0 0 R 0 0 0 0 + Ostracod lining/jaw R + 0 0 + 0 0 0 0 + Taxon diversity 7 7 6 8 11 5 10 6 2 13 Note: +—present; R—rare (<5%), O—occasional (6%–15%), C—common (16%–30%), A—abundant (>30%). Data sources for the sites, which are shown on Figure 1, are as follows. 1—Qarun Lake, Fayoum Depression: Kholief et al. (2007); Hamata mangrove swamp: Kholeif (2007); Suez tidal flat: Kholeif (2004). 2—Nile Delta site 6 and Manzala lagoon site 8: Leroy (1992). 3—Nile cone cores NC-1 and NC-2, Kholeif and Mudie (2009). 4—Aegean Sea sites 3–22 and G5, Marmara sites 2–12, and Black Sea sites 4–13: Mudie et al. (2002, 2004). 5—MAR02-45: Marret et al. (2009). 6—Sites 72* and 1461* of Roman (1974); site 1474 of Wall and Dale (1974). 7—Lake Dzarylgach: Sapelko and Subetto (2007). 8—Lake Saki: Gerasimenko and Subetto (2007). 9—Caspian Sea: Marret et al. (2004). 10—Leroy et al. (2007). 11—Lake Anzaleh (Anz.): Kazanci et al. (2004). 12—Kara-Bogaz Gol (KBG): Leroy et al. (2006). Iran: 13—Nowshahr: Ramezani et al. (2008). 14—Lake Urmia: Djamali et al. (2008); Aral*: Sorrel et al. (2006); Issyk-Kul (IssK*): Giralt et al. (2004). KYRG—Kyrgyzstan.

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Figure 4. Recovery of pollen, dinocysts, acritarchs + prasinophytes, and colonial algae in Black Sea core using acetolysis (samples 72, 1461) versus hydrochloric + hydrofluoric acid (all others) for processing.

systematics of Fensome and Williams (2004) and Marret and Zonneveld (2003), and the fossil fungal spore studies of Kalgutkar and Jansonius (2000) and van Geel (2006). RESULTS Figure 1 shows the sites for which nonpollen palynomorphs have been reported for late Holocene and modern (surface) sediments in the Black Sea corridor, and Table 1 shows the corresponding salinity of the surface water at each site. Because sedimentation rates vary widely among the sites, the table shows all the nonpollen palynomorphs reported for the sites for the historical time interval covering approximately the past 2000 yr. Table 2 is a list of the taxa and their known or presumed biological affinity. The following sections provide an overview of our current knowledge about the biology and ecology of the main nonpollen palynomorphs that have been reported for the surface and latest Holocene sediments the Black Sea corridor. Dinoflagellate Cysts Dinoflagellate cysts are the most common nonpollen palynomorphs throughout the regions of the Black Sea corridor that have been continuously connected to the Mediterranean Sea over the past 2000 yr or longer, but they are absent in the hypersaline intermittently connected or isolated coastal lagoons of the Crimean Peninsula (e.g., Lakes Saki and Dzharylgach). Dinocysts are also absent from late Holocene sediments of both the low- and high-salinity mountain lakes studied in Iran and the brackish lake in Kyrgyzstan (Table 1), although Spiniferites belerius was found in a short interval of penultimate glacial sediment in the presently hypersaline Lake Urmia (Djamali et al., 2008). However, dinocysts are common in the brackish to hypersaline Caspian, Kara-Bogaz Gol, and Aral Seas, where some species

are endemic (Marret et al., 2004). Two taxa (Spiniferites and Lingulodinium machaerophorum) are reported for the low salinity (2‰–4‰) Lake Durankulak-3 (Marinova and Atanassova, 2006), and Spiniferites cruciformis, Impagidinium caspienense, Caspidinium rugosum, and Brigantedinium are present in the oligotrophic freshwater of Lake Sapanca. The most familiar group of dinocysts consists of the spiny gonyaulacoid, autotrophic cysts (Plate 1, figs. 9–10, 13–15, and 17–20) that may be slightly sensitive to oxidation, but usually some are recovered even after acetolysis. With HF processing, however, a much wider range of dinocysts is obtained, particularly including thin-walled “round brown” protoperidinioids (Plate 1, figs. 2–6, 8, 12, and 16) and polykrikoids (Plate 1, fig. 7), which are facultative or obligate heterotrophs (often diatomeating) and flourish in eutrophic water, along with thin-walled toxic gymnodinioid species such as Gymnodinium catenatum (Plate 1, fig. 1). Some gonyaulacoids also proliferate with increased nutrients, including Alexandrium and Lingulodinium. Laboratory experiments of Lewis and Hallett (1997) show that Lingulodinium populations respond more to nutrients and low turbulence than to salinity—they are extremely euryhaline. These authors also reported that cyst populations of the autotroph Lingulodinium machaerophorum respond first to a temperature increase for breaking winter dormancy. Subsequently, the flagellated theca-stage cells need strong stratification for their growth in the photic zone because they are not powerful swimmers. If the photic zone is enriched with nutrients, then red tide population explosions follow, and many resting stage cysts are formed. Therefore, blooms of Lingulodinium cysts may indicate watercolumn stratification and nutrient enrichment more strongly than either salinity or sea level. Principal component analysis of dinocyst core-top data from the Black Sea to Eastern Mediterranean (Mudie et al., 2004) shows two main components: (1) a group of stenohaline halophilic autotrophic species associated with the clear, oligotrophic Aegean Sea waters, and (2) heterotrophic protoperidinioids that are euryhaline and most closely linked with eutrophic surface waters in the Black Sea corridor. It is notable that in this analysis, L. machaerophorum is aligned with the heterotrophic protoperidinioids, suggesting a strong link with high nutrient levels, and explaining the weak correlation of process length and paleosalinity measured by Mudie et al. (2001). Using a larger global data set, however, Mertens et al. (2009) showed that the length and number of processes in L. machaerophorum are also strongly positively correlated with both temperature and salinity at 30 m water depth, possibly related to the sinking rates of the cysts under varying conditions of turbulent mixing versus stratification. Acritarchs and Prasinophytes The term Acritarcha is an artificial category that refers to all small (5–249 µm) microfossils of unknown and varied biological affinities, having a single or multiple-layered organic

Nonpollen palynomorphs

TABLE 2. LIST OF TAXA MENTIONED IN THE TEXT, WITH GENERIC AND SPECIFIC NAMES AND COMPLETE AUTHORSHIPS OF GENERIC AND SPECIFIC NAMES WHERE KNOWN PHYTOPLANKTON SPORES AND PLANT REMAINS Dinoflagellata Acritarcha Alexandrium Halim, 1960 Beringiella Bujak, 1984 Alexandrium monilatum (Howell) F.J.R. Taylor, 1979 Chomotriletes minor (Kedves) Pocock, 1970 Alexandrium tamarense (Lebour, 1925) Balech, 1985 cyst type Chromotriletes Naumova, 1939 Brigantedinium simplex (Wall, 1965) Reid, 1977 = cyst of Chromotriletes rubinus (Christopher, 1976) Fensome et al., 1990 Protoperidinium conicoides (Paulsen) Balech Concentricystes Rossignol, 1962 ex Jansonius and Hills, 1976 Brigantedinium spp. Reid, 1977 ex Lentin et Williams, 1993 = emend. Jiabo, 1978 (misspelled as Concentricystis) (taxonomic Protoperidinium spp. senior synonym = Chromotriletes) Caspidinium rugosum Marret in Marret et al., 2004 Concentricystes rubinus. Rossignol, 1962 ex Jansonius and Hills, Dubridinium caperatum Reid, 1977 1976 Echinidinium transparantum Zonneveld, 1997 Concentricystes rubinus Rossignol, 1962 (partim), p. 134, pl. 2, figs. Gonyaulax apiculata (Penard, 1891) Entz, 1904 5, 6 (nomen nudum) Gymnodinium catenatum Graham, 1943 cyst form Copepod egg-type of Cobricosphaeridum Harland and Sarjeant, Gymnodinium fuscum (Ehrenberg) F. Stein, 1878 1970 emend Head et al., 2003 Gymnodinium mikimotoi Miyake et Kominami ex Oda, 1935 Halodinium Bujak, 1984 Gymnodinium nolleri Ellegaard & Moestrup, 1999 Leiosphaera Eisenack, 1938 (junior synonym for Tasmanites Gymnodinium uberrimum (G.J. Allmann, 1854) Kofoid & Swezey, Newton, 1875) 1921 Leiosphaeridia Eisenack, 1958 Gyrodinium aureolum Hurlburt, 1957 Micrhystridium Deflandre, 1937 Impagidinium caspienense Marret in Marret et al., 2004 Micrhystridium cf. braunii-type of Sorrel et al., 2006 Imagidinium aculeatum (Wall, 1967) Lentin and Williams, 1981 Micrhystridium. cf. M. asagaiense Takahashi, 1964 Impagidinium patulum (Wall, 1967) Stover and Evitt, 1978 Micrhystridium minus Takahashi, 1964 Impagidinium sphaericum (Wall, 1967) Lentin and Williams, 1981 ?Micrhystridium spinuliferum Takahashi, 1964 Impagidinium strialatum (Wall, 1967) Stover and Evitt, 1978 Multiplicisphaeridium Staplin, 1961 Islandinium minutum (Harland et Reid in Harland et al., 1980) Head cf. Multiplicisphaeridium sp. of Struthers, 1996 et al., 2001 Nonpollen palynomorph (NPP) “Type 115” of Pals et al. (1980). Lingulodinium machaerophorum (Deflandre et Cookson, 1955) Pacillina Cleve, 1899 Wall, 1967 = Lingulodinium polyedrum (Stein) J.D. Dodge, 1989 Palaeostomocystis Deflandre, 1937 Lingulodinium machaerophorum forms with short club-shaped Paleostomocystis = Beringiella Bujak, 1984 processes (vars. A–D of Marret et al., 2004) Pseudoschizaea Thiergart and Frantz ex R. Potonié emend. Nematosphaeropsis labyrinthus (Ostenfeld, 1903) Reid, 1974 Christopher, 1976 Operculodinium centrocarpum sensu Wall and Dale, 1966 = cyst of Pseudoschizaea rubina Rossignol ex Christopher, 1976 Protoceratium reticulatum (Claparède et Lachmann) Bütschli, Pseudoschizaea circula (Wolff), emend. Christopher, 1976 1885 Sigmopollis Hedlund, 1965 Operculodinium israelianum (Rossignol, 1962) Wall, 1967 Sigmopollis carbonis Hedlund, 1965 Peridinium Ehrenberg 1830 cyst type Sigmopollis hispidus Hedlund, 1965 = “Type 128” of van Geel, Peridinium ponticum Wall and Dale, 1973 Hallewas & Pals, 1983 Polykrikos Bütschli, 1873 Sigmopollis psilatus Piel, 1971 Polykrikos kofoidii Chatton, 1914 cyst form Sphaeropsis, illustrated by Reid and John (1978) Polysphaeridium zoharyi (Rossignol, 1962) Bujak et al., 1980 = Sporites circulus Wolff, 1934 (partim) Pyrodinium bahamense Tasmanites Newton, 1875 Protoperidinium americanum (Gran et Braarud, 1935) Balech, 1974 Pyxidinopsis psilata Wall and Dale, 1973 Chlorococcales Pyxidinopsis reticulata (McMinn et Sun, 1994) Marret et de Vernal, Botryococcus Kützing 1997 Botryococcus braunii Kützing, 1849 Quinquecuspis Reid, 1976 Pediastrum F.J.F. Meyen, 1829 Quinquecuspis concreta (Reid) Harland, 1977 Pediastrum boryanum (Turpin 1828) Meneghini, 1840 Scrippsiella sensu Head et al., 2006 Pediastrum duplex Meyen, 1929 Selenopemphix nephroides (Benedek, 1972) Bujak in Bujak et al., Pediastrum kawraiskyi Schmidle, 1897 1980, Benedek et Sarjeant, 1981 = cyst of Protoperidinium Pediastrum F.J.F. Meyen, 1829 subinerme (Paulsen) Loeblich III, 1970 Pediastrum simplex Meyen, 1829 Spiniferites Mantell, 1850 Tetraedron Kuetzing, 1845 Spiniferites belerius Reid, 1974 Spiniferites bentorii (Rossignol, 1964) Wall and Dale, 1970 Cynanobacteria Spiniferites cruciformis Wall and Dale, 1973 Anabaena St. Vincent, 1886, ex Bornet and Flahault, 1886 Spiniferites cruciformis morphotype C of Marret in Marret et al., Gloeocapsomorpha Zalesskij, 1918 2004 Gloeotrichia J. Agardh ex Bornet et Flahault, 1887 Spiniferites hyperacanthus (Deflandre et Cookson, 1955) Cookson Rivularia [Roth] Agardh ex Bornet et Flahault, 1887 et Eisenack, 1974 Spiniferites inaequalis Wall and Dale, 1973 Spiniferites mirabilis (Rossignol, 1967) Sarjeant, 1970 Stelladinium stellatum (Wall and Dale, 1968) Reid, 1977 (continued)

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TABLE 2. (continued) PHYTOPLANKTON SPORES AND PLANT REMAINS (continued) Prasinophyta Zygnematales Cymatiosphaera O. Wetzel, 1933 (syn. Pterosperma Pouchet, 1893 Zygnemataceae [Parke et al., 1978] phycoma) Mougeotia C.A. Agardh, 1824 Cymatiosphaera globulosa Takahashi, 1964 (= Pterosperma sp. Spirogyra Link in C.G. Nees, 1820 phycoma) Zygnema C.A. Agardh Hexasterias problematica Cleve, 1900 (syn. Polyasterias Desmidiaceae (sometimes classified in a separate order, problematica Meunier, 1910) Desmidiales) Leiosphaeridia Eisenack, 1958 Cosmarium Ralfs, 1848 Pterosperma Pouchet, 1893 Tasmanites Newton, 1875

Alternaria Nees, 1817 ex Fries, 1821 conidium Callimothalus Dilcher, 1965 type of fruiting body Chaetomium Kunze, 1817 ascospore Coniochaeta ligniaria (Greville) Cooke ascospore Gaeumannomyces Arx & D.L. Olivier; fungal germling similar to hyphopodium of Gaeumannomyces sp. illustrated by Head (1992) Glomus L.R. and C. Tulasne, 1845 chlamydospore type cf. Monosporisporites van der Hammen, 1954 emend. Kalgutkar and Jansonius, 2000

FUNGI Sorosporidium Rudolphi, 1829 type of spore mass illustrated by Kholeif (2004) (syn. Papulosporonites subcircularis Chandra, Saxena & Setty, 1984 type of spore mass) Sporormiella Ellis et Everhart, 1892 Tetraploa Berkeley and Broome, 1850 Thecaphora Fingerh., 1836 Tilletia L.-R. Tulasne and C. Tulasne, 1847 type of teliospores cf. “Type 729” spore of van Geel, Hallewas and Pals, 1983 Valsaria Ces. and de Not. ascospore

ANIMAL REMAINS Protista 1. Protozoa Ciliophora Doflein, 1901 Polychaeta Grube, 1850 Tintinnida Kofoid and Campbell, 1929 “ELENO-2” of van Waveren, 1993 Arthropoda Fusopsis A. Meunier, 1919 lorica type 1. Insecta Pacillina cyst with fimbricated outer wall of Kunz-Pirrung (1998) Chironomidae (nonbiting midges) Tintinnopsis Stein, 1867 “NPP 219” of Van Geel et al. (1989), for the mentum of a Rhizopoda Foraminiferida (planktonic and benthic foraminifera) Chironomid simulid black fly Arcellinida (Arcellaceans): testate amoebae such as Arcella 2. Cladocera Ehrenberg, 1832 Exoskeletal parts (probably the first thoracic antennae, 150 µm long) of a juvenile cladoceran 2. Protozoa Incertae Sedis 3. Crustacea Brunnich, 1772 Radiosperma corbiferum Meunier 1910 (= Sternhaarstatoplast of Branchiopoda Hensen, 1887) Artemia salina (Linnaeus, 1758): brine shrimp Copepoda: copepod egg capsule with fibrous outer wall Rotifera Ostracoda Latreille, 1802 (ostracodes; ostracods; seed shrimp) Filinia longiseta (Ehrenberg, 1834) Leptocythere Sars, 1925 Tardigrada J.A.E. Goeze, 1773 Macrobiotus Schultze, 1834 Chaetognatha (Leuckart, 1854) Hyman, 1959, chaetognaths Macrobiotus hufelandi C.A.S. Schultze, 1834 = egg-type of Jankovska (2007, personal commun.) Platyhelminthes Minot, 1876, flatworms Annelida

Turbellaria Ehrenberg, 1831, flatworms, planarians Neorhabdocoela Meixner, 1938, microturbellarians Note: Further information on the taxonomy and nomenclature of many of the taxa can be found in Fensome et al. (1990).

Plate 1. Light microscope photographs of various ecologically diagnostic dinocysts in the Black Sea corridor. Scale bar is 10 µm. Letters in parentheses are the initials of contributing authors other than P.J. Mudie. Figures 1, 5. Introduced potentially toxic species. (1) Gymnodinium catenatum (F.M.); (5) Alexandrium cyst form (S.E.A.K.). Figures 2–4, 6, 8, 11–12, and 16. Thin-walled cysts of heterotrophic protoperidinioids. (2) Brigantedinium sp.; (3) Peridinium ponticum; (4) Selenopemphix nephroides; (6) Quinquecuspis concreta; (8) Echinidinium transparantum (F.M.); (11) Stelladinium cf. robustum; (12) Quinquecuspis concreta; (16) Quinquecuspis sp. Figure 7. Oxidation sensitive polykrikoid Polykrikos kofoidii (F.M.). Figure 9. Spiniferites cruciformis form 4 (F.M.) first described in the Black Sea early Holocene sediment, but later found common in surface sediment of the Caspian Sea. Figures 10, 17, and 18. Lingulodinium machaerophorum; (10) Caspian Sea morphotype (F.M.); (17) form with clavate spines common in low-salinity waters (F.M.); (18) long-spine form usually only in fully marine environments. Figures 13, 14. Caspian Sea endemics. (13) Caspidinium rugosum (F.M.); (14) Impagidinium caspienense (F.M.). Figure 15. Exclusively freshwater species Gonyaulax apiculata. Figures 19, 20. Exclusively marine species. (19) Impagidinium aculeatum from Nile cone (S.E.A.K.); and (20) Nematosphaeropsis labyrinthus from Aegean Sea.

Nonpollen palynomorphs

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Nonpollen palynomorphs wall surrounding a vesicle with variable ornamentation, and opening by a slit-like or irregular rupture or by a circular pylome (Martin, 1993). A good example of this is found in Sigmopollis (Plate 2, figs. 5 and 6). A comparison of various phenetic classifications of this artificial group is given by Strother (1996). We include the category here because of its long-standing use in geology and because of its use in Traverse’s marine index. However, it is likely that in the future, the term acritarch will be replaced by nonpollen palynomorphs for studies of Pleistocene–Holocene sediments. Prasinophytes and acritarchs, together with dinocysts, are sometimes referred to as organic-walled microphytoplankton (Playford, 2003). However, it has been shown that prasinophyte phycomata are the fossilized cyst-like nonmotile stage of Prasinophyceae (= Micromonadophyceae), which is the oldest group of the green algae (Chlorophyta). The life cycle of some living prasinophytes, e.g., Pterosperma and Pachysphaera, consists of a unicellular motile stage with scale-covered flagella, and a nonmotile floating resting stage called a phycoma (Fig. 5A). The fossil genera Cymatiosphaera (Plate 2, figs. 1 and 2) and Tasmanites (Plate 2, fig. 14) are believed to be the phycomata of Pterosperma and Pachysphaera, respectively (Martin, 1993). A classification of the phylum Prasinophyta is given by Tomas (1993) and by Guy-Olsen (1996), who lists four orders and eight Holocene families of these mainly oceanic planktonic

Plate 2. Light microscope photographs of acritarchs, colonial algae, and zygnemataceans, and scanning electron microscope (SEM) photographs of colonial alga. Scale bar is 10 µm. Letters in parentheses are the initials of contributing authors other than P.J. Mudie. Figures 1, 2. Cymatiosphaera sp., Mackenzie Delta. (1) Midfocus; (2) high focus. Figures 3, 4. Micrhystridium cf. ariakense; Fink Cove coastal pond, Nova Scotia. (3) High focus on dense covering of short spines; (4) mid focus. Figure 5. Sigmopollis sigmoides, Mackenzie Delta; high focus on characteristic sigmoid suture. Figures 6, 8, 9, and 11–13. Black Sea core M02-45, 30 cm. (6) Sigmopollis sp., Black Sea M02-45, 30 cm; (7) Pterospermella, Red Sea (S.E.A.K.); (8) Micrhystridium cf. ariakense, Black Sea; (9) Pseudoschizaea rubina, with irregular polar ornament (details inset); (10) Black Sea core M18 Pseudoschizaea circula, with linear polar ornament. Figures 11–13, 17, and 19, from Mar02-45, SW Black Sea shelf. (11) Spirogyra zygospore; (12) Zygnema-type spore, excysted and partly folded to show projections on inner wall corresponding to surface depressions, 480 cm; (13) Multiplicasphaeridium-type acritarch. Figure 14. Tasmanites, Nile cone (S.E.A.K.). Figure 15. Gloeotrichia filament with terminal heterocyst and lateral akinetospores (P.A. Siver, http://www-cyanosite.bio.purdue.edu/images/lgimages/GLOET1.JPG); (15a–15b) Gloetrichia-type akinetospores from Lake Saki (NG). Figure 16. Anabaena heterocysts from Caspian Sea (S.A.G.L.); (16a) GS18-481-3(20); (16b) CS10237(1). Figure 17. Pediastrum simplex, Egyptian shelf, showing the large size and open structure of most colonies (S.E.A.K.). Figure 18. Hexasterias problematica Black Sea core M18. Figure 19. Botryococcus braunii, Black Sea core Mar02-45, 590 cm. Figure 20. Zygnema spore Lake Ulubat 9/1 (S.A.G.L.). Figure 21. Stenoblast, Hamata mangrove swamp, Red Sea (S.E.A.K.). Figure 22. Radiosperma, SA03K7-300/1 (S.A.G.L.). Figure 23. Gloeotrichia filament, Black Sea Mar02-45. Figure 24. Gloeotrichia filament with endocyst, SA03k71-1 (S.A.G.L.).

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and benthic microalgae. Plate 2 illustrates the main prasinophytes (figs. 1, 2, and 8) and similar nonpollen palynomorphs in the Black Sea corridor cores. Notably, most of these occur in both the brackish to hypersaline Aral Sea (Sorrel et al., 2006), with a fluctuating salinity of ~6‰–15‰ to >100‰ and a history of strong pollution over the past 1000 yr (Aladin and Potts, 1992), and in Lake Issyk-Kul, a slightly brackish (6‰), oligotrophic lake in Kyrgyzstan (Giralt et al., 2004). Modern prasinophytes are regarded as primarily marine, but they also live in brackish and freshwater, and their fossils are associated with lagoonal or deltaic environs. They are most abundant in cold waters, with high nutrients being more important than temperature or salinity, according to Batten (1996). The data for the Black Sea corridor show that Pterosperma (Cymatiosphaera) dominates only

Figure 5. Life cycles of algal nonpollen palynomorphs: (A) Pterosperma (from Martin, 1993); and (B) Pediastrum boryanum (from Batten, 1996). Stages: 1—coenobium; 2–5—asexual reproduction stages: 2, 3—release of vesicle with zooids, 4—new coenobium forming within vesicle, 5—release of coenobium, 5f–5k—sexual reproduction stages.

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in the lower-salinity eutrophic waters, and it is absent in the oligotrophic marine–hypersaline Aegean waters. In contrast, Pterospermella (Plate 1, fig. 7) is only reported for the hypersaline Red Sea, and the small spiny acritarch Micrhystridium (Plate 2, figs. 3, 4, and 8) was found only in the Red Sea, the hypersaline Crimean liman-lagoons, and, as Micrhystridium braunii–type, in the Aral Sea. The acritarch Polyasterias problematica is found in the Caspian and Aral Seas (as Hexasterias problematica; Sorrel et al., 2006); it is usually considered as a brackish to marine organism, but appears to be euryhaline (Matthiessen et al., 2000; Sorrel et al., 2006). The acritarch Radiosperma corbiferum was present mostly in the brackish intervals of the Aral Sea (rare and possibly reworked in the hypersaline intervals) and in the modern sediments of the freshwater Lake Sapanca east of the Marmara Sea (Leroy and Albay, 2010). The biological affinity of this nonpollen palynomorphs is unknown; a morphologically similar organism in the plankton of the Gulf of St. Lawrence, Canada (Bérard-Therriault et al., 1999, their plate 149b), is listed as Protiste sp. 2, in the group Protiste incertains (Uncertain Protists). Brenner (2001) reported that an apparently similar morphotype Radiosperma cf. corbiferum has a widespread distribution ranging from low-salinity Arctic estuaries to upwelling marine waters off Peru. Colonial Algae The colonial algae are unicellular Chlorophyceans that form colonies (coenobia) from a number of cells linked together and arranged in a specific pattern. In Pediastrum (Plate 2, figs. 17 and 18), the colonies have a fixed number of cells, the coenobia are flattened, and reproduction is by means of biflagellated cells, either aggregated within a vesicle that is liberated from the parent colony cell, or during sexual reproduction, and they are briefly free-swimming (Fig. 5B). In Botryococcus, the colonies range from small (~30 µm) subspherical clumps of cells (Plate 2, fig. 19), embedded in an oil-rich matrix, to relatively large (~140 µm) grape- or mulberry-like (botryoidal) groupings held together by gelatinous fibers or membranes (Batten and Grenfell, 1996). In the Black Sea corridor, the distribution of the colonial algae Pediastrum and Botryococcus shows the same trend as for Pterosperma (Cymatiosphaera). These nonpollen palynomorphs are abundant in parts of the Black Sea, sparse in the Marmara Sea, and almost absent in the Aegean and Levantine Seas, but they reappear in the Nile Delta and on the Egyptian shelf. These colonial algae are commonly considered to be indicators of riverwater inflow (e.g., Head, 1992, and references therein). However, they also appear in sediments of the Aral Sea and the hypersaline Kara-Bogaz Gol of the Caspian Sea, and Botryococcus braunii ha been found in plankton surveys of the Aral Sea (Piontkovski and Elmuratov, 2008). Detailed studies of surface sediment samples in the Baltic Sea show that Pediastrum boryanum and P. kawraisky are dominant in salinities from 6‰ to 8‰ (full range is 5‰–9‰), while P. simplex and P. duplex occur in salinities of less than 3‰–5‰ (Matthiessen and Brenner, 1996). The Baltic

study also shows that Botryococcus cf. braunii tolerates salinities up to 8‰. Zalessky (1926) reported that large botryoidal colonies of B. braunii occurred in the freshwater Russian Lake Beloë, while those in the brackish water (4‰) of Lake Balkash (Kyrgyzstan) were closely packed and globular. Growth of Botryococcus may be favored by seasonally cold, oligohaline conditions (Batten and Grenfell, 1996), in either eutrophic shallow or deep oligotrophic lakes (Chmura et al., 2006); in contrast, blooms of Pediastrum in the Canadian Great Lakes are triggered by excess phosphate loading (Nicholls, 1997). Zygnematales The Zygnematales are charophycean green algae that reproduce by conjugation to produce resting spores or zygotes with a sporopollenin-like cell wall. This order includes unbranched filamentous green algae of the families Zygnemataceae and singlecelled Desmidiaceae (sometimes classified in a separate order Desmidiales). Modern Zygnemataceae typically live in shallow, stagnant freshwater lakes, ponds, or in wet soil (Van Geel, 2001) and produce spores in the spring when conditions are warm. In late Holocene sediments of the Black Sea corridor, zygospores known to be formed by Zygnematacean algae have been found only in low-salinity lakes or freshwater lakes. These include Spirogyra, which was found in Lake Durankulak (Marinova and Atanossova, 2006), the montane inland Lake Izzyk-Kul, Lake Sapanca, and Lake Manzala of the Nile Delta. Zygospores of Spirogyra and the Zygnematacean genera Mougetia and Debarya were also found in sediments of the Nile Delta (Leroy, 1992). Zygnema spores (Plate 2, fig. 20) are common in the polluted, freshwater Lake Ulubat south of the Marmara Sea (Fig. 1). Pseudoschizaea rubina (Plate 2, fig. 9) may be a zygospore of the zygnematacean alga Debarya (Grenfell, 1995), although this relationship has not been confirmed by laboratory cultures, and it is often simply grouped with the acritarchs, as in Tables 1 and 2. This sphaeromorphitic acritarch, with distinctive concentric markings on both hemispheres of its dorsoventrally flattened vesicle, was first described as Sporites circulus in Pliocene brown coals, and then as Concentricystes rubina in marine sediments off Israel. Pseudoschizaea rubina is distinguished from the similar species Pseudoschizaea circula by an irregular, maze-like polar complex up to one quarter of the vesicle diameter (Christopher, 1976). Concentricystes s.l. is usually considered to be a freshwater alga because of its association with wadi or river terrace deposits (Christopher, 1976), and some species, including P. circula, have only been recorded from terrestrial or fluvial environments. However, both P. circula and P. rubina are occasionally present in the late Holocene marine sediments of the Black Sea but are absent from the early Holocene brackish-water sediments. Pseudoschizaea circula is rare in the hypersaline Lake Saki, and unspecified Pseudoschizaea species have been found in the Nile Delta, Red Sea, and Aral Sea. Desmids are most common in oligotrophic freshwater lakes and ponds, but some species e.g., Closterium aciculare,

Nonpollen palynomorphs mark eutrophic conditions (Graham and Wilcox, 2000). Desmid zygotes Coelastrum and Mougeotia occur in Lake Durankulak (Marinova and Atanassova, 2006), but in the Black Sea, Mougeotia and Closterium have been reported only for mid-Holocene sediments (Mudie et al., 2010), although they are markers of river transport to modern sediments in the Beaufort Sea (Matthiessen et al., 2000). Zygotes of Tetraedon and Coelastrum are sporadically abundant in the freshwater Lake Sapanca. Cyanobacteria Fossil Cyanobacteria (blue-green algae) are rare in the Black Sea corridor sediments, although the marine unicellular Synechococcus cynobacteria occurs in both eutrophic and oligotrophic waters of the Black Sea corridor (Uysal, 2006). Gloeotrichia is rare to common in the modern sediments of Lake Sapanca (Table 1; Plate 2, fig. 24), but in the Black Sea, Gloeotrichia-type sheaths were found only in the early Holocene sediments (Plate 2, fig. 23). Van Geel (2001) noted that Gloeotrichia marks nutrient-poor conditions in late glacial lakes because it is a nitrogen-fixing alga that subsequently makes conditions suitable for other aquatic plants. The planktonic filamentous heterocyst and akinete-producing alga Anabaena is present in both IssykKul and the Caspian Sea (Plate 2, fig. 16). Fungi and Animal Remains Other important nonpollen palynomorphs in the Black Sea are derived from fungi and various types of planktonic or benthic zooplankton. Plate 3 illustrates various fungal spores, conidia (Plate 3, figs. 1–5), ascomata (Plate 3, fig. 6), or germlings from Black Sea core 45. There are very few marine fungi, and most of the spores and other fungal remains must have been transported to the inland seas by air or soil erosion. Some dung fungi, such as Sporormiella (Plate 3, fig. 29), mark nutrient enrichment from domestic animals and are a reliable proxy for faunal biomass (van Geel and Aptroot, 2006). Several spore types (e.g., Tilletia, Ustilago) are produced by parasites of specific native plants and domestic crops. In archaeological middens, Glomus-type fungal spores (Plate 3, figs. 24 and 28) are extremely resistant to fire and biological degradation; hence, they persist in shell middens and soils after almost all the pollen has disappeared from degradation (Bryant and Holloway, 1983; Leroy et al., 2009), and they may survive very long-distance transport by river water. Glomus-type spores are present on the Nile cone and the Red Sea, in addition to the Black Sea and almost all lakes, unless the catchment is very small and the sample is far from the shore. Various zooplankton remains have been reported for the Black Sea corridor, including the chitinous skeletal remains of a juvenile cladoceran (Plate 3, fig. 21), copepod egg capsules (Plate 3, fig. 19) and their eggs (Plate 3, fig. 20), the organic linings of benthic microforaminifera (Plate 3, figs. 7, 13, 14), polychaete worm jaws and pincers (known to geologists as scolecodonts; Plate 3, fig. 10), and various morphologically similar

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palynomorphs that include arthropod sclerites (Plate 3, fig. 15). The jaws of unknown ostracods were identified in Lakes Anzaleh and Sapanca (Plate 3, fig. 26) and in early Holocene sediments of the Black Sea (Plate 3, fig. 22). The lining of a very small ostracod (Plate 3, fig. 18; probably Leptocythere according to David Horne, 2007, personal commun.) is found in the Black Sea, and ostracod linings were also present in Lake Saki and on the Egyptian shelf. Eggs of the rotiferan Filinia longiseta were found only in the freshwater Lake Sapanca. Several types of tintinnid loricas are reported for the Aral Sea (Sorrel et al., 2006) and the Black Sea (Plate 3, fig. 8), and a Tintinnopsis is present in Lake Sapanca (Plate 3, fig. 9). The marine palynomorph Palaeostomocystis (Plate 3, fig. 11) also resembles tintinnid loricas found in arctic regions (Matthiessen et al, 2000); it is rare in the Marmara Sea. Other relatively large (~60–80 µm), brown, vase-shaped palynomorphs (Plate 3, figs. 16 and 17) occur in the Marmara Sea and on the Egyptian shelf. These nonpollen palynomorphs resemble the resting eggs (oocytes) or egg capsules/cocoons of microturbellarian flatworms (Platyhelminthes, Order Neorhabdocoela), which are mostly freshwater organisms (Haas, 1996); however, other turbellarians are common in coastal marine environments, and predatory marine flatworms parasitize mussels in the Black Sea (Murina and Grintsov, 1998). According to Ole Bennicke (2008, personal commun.), the egg walls of the marine flatworms tend to be thicker than those of the freshwater taxa. Morphologically similar nonpollen palynomorphs that occur in deep-sea sediments of the Banda Sea have been referred to the chitinous loricas of marine tintinnids (van Waveren, 1994). Clearly, more research on this group of palynomorphs is required before these nonpollen palynomorphs can be used reliably as environmental indicators when recovered from brackish-water or marine sediments. There is also further need to confirm the biological link between the brackish-water nonpollen palynomorphs referred to the tintinnids: several illustrations of vase- or urn-shaped testate amoebae living in peatland moss (Swindles and Roe, 2007) are very similar in size and morphology to some of the tintinnid nonpollen palynomorphs found in the Black Sea corridor. The organic-walled palynomorph Halodinium (Plate 3, fig. 12) was first recorded and described as an acritarch of unknown affinity occurring in subarctic marine sediments of the Bering Sea, and it is widely distributed in the Arctic (Matthiessen et al., 2000), including ponds of the Mackenzie Delta. The shape of Halodinium is similar to that of the testate amoebas Cyclopyxis and Arcella (illustrated by Beyens and Meisterfeld, 2001), and it is possible that this palynomorph is the organic lining of a testate amoeba (thecamoebian). The testate amoebas are a polyphyletic group of protozoans, the largest group (75%) being the Arcellinida. The empty tests remain intact after death of the amoeba and can be recovered fully from anaerobic sediments by dispersion in water and gentle sieving, but they decompose within a few weeks under aerobic conditions (Beyens and Meisterfeld, 2001). Extraction by palynological processing with acids and/or alkalis, however, gives variable recovery and does not produce a reliable picture of the thanatocoenoses (Swindles and Roe, 2007).

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Nonpollen palynomorphs Halodinium is rare in the Black Sea and present in the Nile Delta and Red Sea. Unspecified thecamoebians have been reported for the calcareous mire near Nowshahr in northern Iran and are present in some freshwater lakes of the region. The distribution of these nonpollen palynomorphs in the surface samples of the Black Sea corridor (Table 1) shows that the fungi are most abundant in the low-salinity waters of the inland lakes, and the Black and Marmara Seas, they disappear in the Aegean, and then they return off the Nile Delta, being clearly linked to the proximity of terrestrial environments where they originate. In contrast, microforaminiferal linings continue to be present in the marine and high-salinity waters of the northern Aegean Sea (Aksu et al., 1999; Kotthof et al., 2008) and the Egyptian shelf (Kholeif, 2010), but they are absent from both the anoxic deep basins of the southern Black Sea and from late Holocene marine sediment of the Nile cone in the hyperoligotrophic water of the Levantine basin. Microforaminiferal linings might be expected to be good markers of sustained marine connection with the Mediterranean, but they also occur in the Kara-Bogaz Gol, which, during most of the late Holocene, was only connected to the global oceans by a canal linking the Volga River and the Azov Sea. They are also present in the hypersaline water of Lake Dzharylgach, which has been isolated from the Black Sea during historical time. Kotthoff et al. (2008) reported

Plate 3. Light microscope photographs of fungi and animal remains from Black Sea cores M02-45T, 30 cm, Marmara Sea core M98-12, Caspian Sea core 31, Lake Sapanca samples (SAK), Lake Dzharylgach, Egypt (NC), and the Red Sea. Scale bar 10 µm. Letters in parentheses are the initials of contributing authors other than P.J. Mudie. Figures 1–5. Fungal spores from SW Black Sea. (1, 2) Tilletia-type teliospores, M45T; (3) Chaetomium ascospore, M45T; (4) Valsaria sp. Ascospore; (5) Coniochaeta ligniaria ascospore (MM-F). Figure 6. Fungal germling similar to hyphopodium of Gaeumannomyces sp., B7, 70 cm. Figure 7. Caspian microforaminiferal lining CS31-3/2 (S.A.G.L.). Figure 8. Pacillina-type tintinnid lorica. Figure 9. Tintinnid lorica, with embedded charcoal fragments, SA0361 (S.A.G.L). Figure 10a. Simple scolecodont, core NC-1, Egyptian shelf (S.E.A.K.). Figure 10b. Bifurcated scolecodont; Red Sea (S.E.A.K.). Figure 11. Paleostomocytis sp.; Figure 12. Thecamoeban cf. Halodinium minor. Figures 13, 14, Microforaminiferal linings; (13) trochospiral form, NC-1, Egyptian shelf (S.E.A.K.); (14) Black Sea Mar02-45 planispiral, open lining. Figure 15. Arthropod sclerite, Lake Ulubat, AK104_24.5_1 (S.A.G.L.). Figures 16, 17. Cf. tintinnid loricae/turbellarian egg capsules from core M98-12, 30 cm; (16) lorica type 1, with short apiculate base; (17) lorica type 2, with rounded base. Figure 18. Organic lining of brackish water ostracod, cf. Leptocythere, M02-45, 790 cm. Figures 19, 20. Copepod eggs; (19) copepod eggs; (12) ?copepod egg capsule with fibrous outer wall; M45. Figure 20. Spiny copepod egg. Figure 21. Exoskeletal parts (probably the first thoracic antenna, 150 μm long) of a juvenile cladoceran, M45P, 440 cm. Figure 22. Mouth part of a small ostracod. Figure 23. Sorosporium-type spore mass; Red Sea (S.E.A.K.). Figure 24. Glomus-type fungal spore, NC core 2 (S.E.A.K.). Figure 25. M45T Desmid or Tardigrade egg. Figure 26. Ostracod jaw CP14, 45 cm (S.A.G.L.). Figure 27. Unknown amphipod from Lake Dzharylgach, Dz70, 160–155 cm (T.S.). Figure 28. Glomus group SA03K71-13 (S.A.G.L.). Figure 29. Sporormiella spore SAK71-1 (S.A.G.L.).

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that microforaminiferal linings are among the most oxidationsensitive palynomorphs. These palynomorphs are abundant and well-preserved at ~1000 m water depth in the Mt. Athos Basin of the northwest Aegean; therefore, their absence in the anoxic ~2-km-deep basins of the southeastern Black Sea strongly suggests that few benthic foraminifera live under conditions of low surface salinity (<22‰) and anoxia below 200–400 m water depth. Planktonic foraminifera are present in the Mediterranean and Marmara Sea but absent from the Black Sea and the isolated inland seas. DISCUSSION At this stage of our study, it is not possible to define many clear patterns in nonpollen palynomorphs distributions within the Black Sea corridor, in part because of the different processing methods used, and also because many phytoplankton taxa will grow under a very wide range of temperature and salinity conditions, although they only bloom under optimal conditions. Presence or absence data are therefore not very diagnostic except for taxa with very restricted distributions, such as the brackish-water (2‰–13‰) dinocyst species Impagidinium caspienense, Caspidinium rugosum, and Spiniferites cruciformis. Recently, small amounts of I. caspienense, S. cruciformis, and Brigantedinium spp. were found in the modern (0–55 yr) sediments of freshwater Lake Sapanca between the Marmara Sea and Sakarya River, but it is presently unknown if they are in place or a recent introduction that may not survive in the freshwater environment (Leroy et al., 2009). Other well-known Mediterranean dinocysts are notably absent from the semimarine water of the Marmara Sea, the low-salinity marine waters of the Black Sea, and inland seas or lakes; these taxa include Impagidinium aculeatum, I. patulum, I. sphaericum and I. strialatum, Nematosphaeropsis labyrinthus, Operculodinium israelianum, and Polysphaeridium zoharyi. So far, no dinocysts have been found in either the hypersaline limanlagoon sediments of the Crimean Peninsula that were treated with hot HCl, or the saline sediments of the Egyptian Fayoum basin (Kholeif et al., 2007), where the normal HCl-HF method of processing was used for palynomorph processing. The usual methods used for palynological processing and the examination of residues mounted on microscope slides are not the best techniques for study of the animal remains of Ostracoda, Cladocera, and Chironomidae because only small parts of the whole fossil are recovered. Other techniques exist that are specific to each arthropod group and can provide a wealth of new data on paleotemperature, paleosalinity, and paleonutrients. Likewise, thecate amoebians are fully studied using different extraction and observation methods than those used in palynology. The recovery of just a few organic linings and sclerotized or chitinized body parts, however, may provide valuable clues to the need for further specialized studies of sedimentary intervals where the skeletons of animals are rare because of unsuitable preservation, including dissolution of carbonate and silicate shells. For example, in the early Holocene sediments of the Black

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Sea, foraminiferal and ostracod linings were recovered in older sediments below those in which the animals were first visible in sieved samples examined by binocular microscope. In the Black Sea corridor, the most restricted distributions are shown by an unnamed desmid/tardigrade egg (Plate 3, fig. 25) and acanthomorphic acritarch (Plate 2, fig. 13) so far reported only for the Black Sea. Halodinium (Plate 3, fig. 12) and Paleostomocystis may also be confined to the Black Sea. Possibly, these nonpollen palynomorphs can help us distinguish the paleoenvironments of inland seas with quasi-permanent interglacial marine connections from the inland lakes that have been isolated throughout the Holocene. Rotifer eggs (including those of Filinia) were found only in Lakes Anzaleh and Sapanca, and the tintinnid genus Tintinnopsis was found only in Lake Sapanca. Nonpollen palynomorphs identified as the mouthpart or jaws of polychaetes so far have been reported only for the Marmara and Red Seas, suggesting that they indicate a neritic marine influence, as noted by Head (1992). In contrast, the mandibles of ostracods have been found only in the low-salinity or freshwater lakes, and an unnamed nonpollen palynomorph that may be an amphipod mouthpart (Plate 3, fig. 27) so far has been found only in the hypersaline liman-lagoon of Lake Dzharylgach. Other large (>200 µm) arthropod parts from lake sediments include a sclerite from an aquatic Hemipteran (Plate 3, fig. 15), and it is possible that the smaller (<100 µm) rod- or Y-shaped scolecodonts resembling simple polychaete worm jaws may actually include sclerotinized rostral spines of freshwater or marine copepods (van Waveren, 1994) or the claws of chironomid larvae (Plate 3, fig. 10B; according to Dirk Verschuren, 2008, personal commun.).Throughout the Black Sea corridor, there is a conspicuous absence of the gelatinous filaments of the periphytic Rivulariaceae (sometimes classified with the Cyanobacteriales), which includes marine littoral and freshwater species, most commonly in clear, unpolluted, streams, but also found in stagnant waters (Graham and Wilcox, 2000). The hypersaline Lake Saki (~45°2′N, 33.5°E, water depth <1 m) is of considerable interest because it is a possible analogue for an early Holocene Black Sea arid environment (Gerasimenko and Subetto, 2007). At present, Lake Saki is a salt lake (predominantly NaCl) separated from the Black Sea by a large sand bar up to 5 m high. The liman-lagoon occupies the estuary of a former drainage basin that was cut off by the construction of two dams in 1895 (Solomina et al., 2005), and its freshwater inflow is now limited to precipitation and Black Sea water. This liman-lagoon has a salinity range of 40‰–80‰ in its outer “natural” part and 15‰–200‰ in the inner dammed section. The liman-lagoon contains black to gray clay annually varved sediments: eolian sediments accumulate mainly in summer, and total varve thickness is a proxy for intensity of river erosion, which corresponds to the precipitation regime. Time-series analysis of the varves and local pine tree rings show that both record regional moisture regimes that are correlated with shifts of the North Atlantic Oscillation and its East Atlantic–Western Russia teleconnection (Solomina et al., 2005). It is of major interest to find that both

sectors of the alkaline Lake Saki sediments have high concentrations of well-preserved pollen, because, in some circumstances, these alkaline environments are problematic for pollen preservation (Bryant and Holloway, 1983). In other situations, such as laminated sediments of the Dead Sea, however, the preservation is excellent (Heim et al., 1997), and it is possible that the poor preservation is more linked to oxidation than pH. In Lake Saki, nonpollen palynomorphs are extremely sparse in the inner limanlagoon, but low-diversity assemblages are common in the outer section. The nonpollen palynomorph assemblages are dominated by Micrhystridium and colonial algae. Well-known cyanobacterial palynomorphs are unexpectedly absent, given that the filamentous algae Cladophora and Gloeotrichia grow in the lake (Mukhanov et al., 2004). Cladophora is not known to produce fossilizable resting spores, but elsewhere in the Black Sea corridor, heterocysts of Gloeotrichia are found (Plate 2, fig. 16). In Lake Saki, it is possible that very common unicellular brownish spores (Plate 2, figs. 15a and 15b) are akinetes of Gloeotrichia. Fungal spores are also common in Lake Saki, including Glomustype spores that are derived from mycorrhizal root-fungi and thus indicators of severe soil erosion. Pseudoschizaea cf. S. circula is rare. Overall, the Lake Saki nonpollen palynomorph assemblage bears no resemblance whatsoever to the early Holocene nonpollen palynomorph assemblages in the Marmara and Black Seas, and contradicts the idea that the outer Black Sea shelf was aerially exposed and dissected by estuaries with sand barriers and liman-lagoons at that time. Lake Dzharylgach (~45°34.7′N, ~32°51.7′E) is similar in size and water depth to Lake Saki, but it was progressively disconnected from the Black Sea from ca. 4.7 to 2.57 ka, and has been an isolated salt lake for more than 2000 yr, with a presentday salinity of ~100‰. Like the inner part of Lake Saki, pollen and spores are common and nonpollen palynomorphs are sparse in Lake Dzharylgach, and include only Micrhystridium, Charophycean remains, rare fungal spores (not the Glomus-type), linings of ostracods, and planispiral open microforaminifera. Pediastrum is rare, being present only in the interval just before 2000 yr B.P. The deep time (>14,000 yr) distribution of some nonpollen palynomorphs in the Marmara Sea core 98-12 (Mudie et al., 2002) shows that that there are no very clear patterns of palynomorph increase/decrease except in the pollen and total organic carbon (TOC) signals after 10 ka, and in the peak of multicellular fungal spores that coincides with the high sediment influx ca. 14–10 ka. The Pediastrum, prasinophyte (mostly Pterosperma), and acritarch records (mostly Sigmopollis) show variable oscillations that are not clearly linked with lithofacies or cyclicity. In the southwestern Black Sea core MAR02-45 (Marret et al., 2009; Hiscott et al., 2007), where sample processing without acetolysis yielded a full complement of both oxidation-resistant and oxidation-sensitive peridinioid dinocyst heterotrophs (Fig. 6), we find a succession from assemblages dominated by the brackish-water species Spiniferites cruciformis and

Figure 6. Distribution of nonpollen palynomorphs in Black Sea core M02-45 (from Marret et al., 2009). Alpha-2 marks a regionwide sedimentary hiatus.

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Pyxidinopsis psilata that live in the Caspian today, to a gradual replacement by the red tide–producing, euryhaline, eutrophic species Lingulodinium machaerophorum and protoperidinioid heterotrophs that thrive in highly eutrophic waters. The colonial alga Pediastrum only dominates from ca. 9.3 to 8.6 ka, during the gradual sea-level rise and salinity increase from ~5‰–13‰ to 16‰, and then it becomes sparse in the upper Holocene, where salinity is 16‰–20‰, like today. Zygnematacean spores of Spirogyra and Zygnema (Plate 2, figs. 11, 12, and 20) are confined to this interval. A few semicells of Cosmarium were also found in the overlying sapropelic silty mud unit, together with occasional Mougeotia spores. Few Cosmarium zygotes were observed, suggesting that the desmids were not growing and reproducing in the sea but were transported as empty hemicells from onshore habitats. As noted for the St. Erth nonpollen palynomorph assemblages (Head, 1992), it is likely that most of the sporadically occurring Zygnematacean spores in the Black Sea are transported by rivers. There is one small peak of Botryococcus at the end of the initial period of Mediterranean inflow of sulfur-rich water (ca. 8.6–8 ka), which also marks the onset of strongly stratified conditions that would favor Lingulodinium. Linings of the brackish-water ostracods overlap with the peak of the combined colonial algae coenobia, and then the microforaminiferal linings become more prevalent. There is a notable absence of the acritarch Radiosperma corbiferum or Radiosperma cf. corbiferum, which, together with Operculodinium centrocarpum, characterizes the slightly brackish Yoldia Sea phase of the Baltic Sea (Brenner, 2005, 2006). For future studies, it would be helpful to describe and illustrate the microforaminiferal linings found in the Caspian/Aral Seas because distinctive grouping of benthic foraminifera, e.g., elongate uni- or biserial miliolids versus spherical trochamminids (Fig. 7), can be distinguished by the similar shape of their chitinous linings: uni-/biserial, planispiral proximate/open, trochospiral, etc. (Mathison and Chmura, 1995; Stancliffe, 1996). It is possible that some of these nonpollen palynomorphs are the key markers for the inland seas that have been essentially isolated from the world’s oceans for the past 15,000 yr, compared to the more persistently connected Black and Marmara Seas. It may also be useful to record the species of Pseudoschizaea (Fig. 7) because P. rubina seems to have a more limited distribution in the coastal Eastern Mediterranean region, whereas P. circula has a very widespread distribution, occurring from mountain to coastal regions. More information is also needed about ostracod versus polychaete jaw characteristics. Saniawski (1996) illustrated the chitinous mandibles and jaw supports of two extant polychaete worm families, but there is little information on Holocene taxa, and at present, it is not clear how to distinguish small simple polychaete worm jaws from the sclerotized mouthparts of either Cladocerans (Korhola and Rautio, 2002) or Chaetognaths (arrow worms). The toothed jaws of scolecodonts from the Marmara Sea are similar to Scolecodont Form 4 of Head (1993, fig. 6-4), whereas those found in the mangrove swamp sediment of the Red Sea and Egyptian shelf are simple (rod-shaped) or bifur-

cated and more similar in shape to sclerites from body parts of small arthropods. Similar bifurcated jaws are found in beach and shallow-marine deposits of southwestern India (Limaye et al., 2007). Relatively large (>100 µm) sclerites of arthropods are frequently found in freshwater lakes and lagoons, e.g., Lakes Ulubat and Sapanca, and in the inland brackish Caspian Sea. The size and shape of these nonpollen palynomorphs differ considerably from the sclerotic mouthparts of copepods and polychaetes found in continental shelf sediments off California (Mudie, 2009, personal observation). Ostracoda, Cladocera, and Chironomidae are arthropod groups that have received a lot of attention recently in Holocene studies owing to their potential to reconstruct past water temperatures and salinities (Smol et al., 2001). Ostracod mandibles (Fig. 5) and shells (Plate 3, fig. 18) and Cladoceran (Plate 3, fig. 21) remains are complex structures, and their identification is outside of the normal specialization of a palynologist. The jaws of ostracods (Plate 3, figs. 22 and 26) can be distinguished from the similar but more complicated mouthparts of chironomid midge larvae (Fig. 5; Walker, 2001; Eggermont et al., 2008). Chironomids have a joined symmetrical left and right part with identical rows of teeth, and the palynomorphs are usually found with cojoined central and ventral jaw parts. In contrast, the left and right teeth of ostracods are not joined, and the palynomorphs are found as single jaws with toothed ends. Anthropogenic influences in the Black Sea corridor are most clearly marked by changes in forest pollen influxes (Mudie et al., 2007) that correspond to deforestation after ca. 7.5 ka and by the appearance of cereal pollen around 5 ka, followed by various horticultural taxa, e.g., olives and vine grapes. However, nonpollen palynomorphs are unique palynomorphs for tracing the history of livestock production from marine sediment records. For example, the occurrence of herbivore dung fungal spores of Sporormiella at coastal sites off Varna clearly shows that livestock agriculture was practiced by the Bronze Age, but there is no evidence of earlier animal farms. In Lake Durankulak, dung-fungus spores of Chaetomium, Coniochaeta, Podospora, and Sordaria indicate extensive local stockbreeding and grazing during the Early Bronze Age (Marinova and Atanassova, 2006). Peaks of Neurospora spores, a fire indicator, correspond to maxima in charcoal particles and Glomustype spores that indicate soil erosion (Marinova and Atanassova, 2006). The smut fungus Sorosporium parasitizes mainly grasses, including important crops like sorghum, maize, and millet, and the presence of Sorosporium-type spores in the Red Sea marks long-distance transport from highland regions where the crops are mainly grown (Kholeif, 2004). In Issyk-Kul, increased Botryococcus and fungal spore percentages are taken as signs of greater erosion and heavier grazing in the catchment after A.D. 1560 (Giralt et al., 2004). Dinocyst records are also important for understanding the history of toxic red-tide blooms in the Black Sea corridor. In MAR02-45, it is clear that blooms of the toxic species L. polyedrum are endemic to the Black Sea waters. In contrast, the late

Nonpollen palynomorphs arrivals of the toxic red-tide species Gymnodinium catenatum and Alexandrium-cyst species around 2400 and 500 yr B.P., respectively, appear to be related to recent introductions via ship-ballast discharge, which today persists as a serious environmental problem in the Black Sea (e.g., Moncheva and Kamburska, 2002). Recent toxic cyst–forming species include Alexandrium monolatum (first seen in 1991), Gymnodinium uberrimum (1994), G. fuscum (1970–1980), and Gyrodinium cf. aureolatum or G. mikkimotoi (1970–1980). These cysts have not yet been recovered as fossils in the surface sediments, however, so it is not yet clear if they survive and reproduce in the Black Sea.

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CONCLUSIONS There is a large range of nonpollen palynomorphs in the Black Sea corridor that remain to be fully exploited by palynologists. Only a small proportion has been identified and related to living organisms or, failing that, have been given a name or a type number. Nonetheless, the change from use of geological fossil names (e.g., Cymatiosphaera) to biological names (e.g., Pterosperma) by Quaternary paleoecologists illustrates the great progress made in the study of Holocene micro- and macrofossil groups during the last decades. Within much of the old geological

Figure 7. Diagrams of characteristic features of nonpollen palynomorphs with possible environmental importance in the Black Sea–Mediterranean corridor. (A-1) Ostracod mandible, showing last mandible coxa and teeth (from Horne et al., 2002); scale bar = 50 µm. (A-2) Chironomid mouthparts (from Walker, 2001). m—mentum; v—ventromentum plates. (B) Microforaminiferal shapes (after Stancliffe, 1996): 1—single chamber; 2—uniserial;3—biserial type ii; 4—coiled biserial; 5—coiled uniserial; 6—planispiral; 7—spiral uniserial; 8—spiral coiled; 9—trochospiral. (C) Pseudoschizaea species (after Christopher, 1976).

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information, the new biological identifications have led to new, more in-depth interpretations of environmental change. These powerful new paleoenvironmental tools are likely to spread fairly quickly to the whole of the Pleistocene and further down the Cenozoic. Depending on the awareness of the various palynologists, and the use of acetolysis versus hydrofluoric acid for sediment processing, only a small portion of the nonpollen palynomorphs in marine/brackish water lakes in the Black Sea corridor has been regularly counted and published. However, our overview of the available, mostly qualitative data clearly shows the value of extracting the full complement of nonpollen palynomorphs, particularly thin-walled peridinioid dinocysts and fungal spores, because of the potential for detailed interpretation of paleoenvironmental conditions and human activity (e.g., livestock husbandry, agriculture, and burning). The most important initial results are summarized as follows. 1. The surface distributions of the nonpollen palynomorphs dinocysts, Pediastrum and Botryococcus, zygnematacean algae, and zooplankton remains show that during the early Holocene, the Black Sea was a brackish sea like the modern Caspian Sea or the outer Baltic Sea: It was not a freshwater lake. The presence of Pediastrum may indicate relatively high phosphorus levels from river inflow, but there are no nonpollen palynomorph indices of fire, soil erosion, or human settlement. 2. Core-top samples from the wide range of salinity in the Black Sea corridor show that the prasinophyte Pterosperma (fossil name Cymatiosphaera) and other small unicellular acritarchs are usually more abundant in low-salinity environments and they include freshwater species; hence, Traverse’s marine influence index cannot be used as a reliable marker of detailed sea-level change. 3. Peaks in the fossil Cymatiosphaera globosa (= Prasinophyte Pterosperma), colonial algae, and some dinocysts, e.g., Lingulodinium machaerophorum (= L. polyedrum), mainly reflect nutrient levels and stratification of the water column, and are not reliable markers of sea-level change. 4. Fungal spores appear to be the best index of terrigenous input from soil erosion, and they are important markers of Bronze Age farming practices. 5. Laminated sediments in the hypersaline liman-lagoons of the Crimean Peninsula are characterized by high concentrations of pollen regardless of salinity, by an absence of dinocysts, and variable amounts of low-diversity nonpollen palynomorph assemblages. Ponds with NaCl salt concentrations greater than 100‰ have few nonpollen palynomorphs compared to ponds with a salinity of 40‰–80‰. High concentrations of the acritarch Micrhystridium may characterize these environments. 6. In palynological preparations for arthropod and polychaete groups, only partial information is left that does not allow detailed paleoenvironmental reconstruction compared to the algal nonpollen palynomorphs groups, e.g., Chlorophyceae and Cyanobacteria, where identification to generic and species level is possible and leads to significant enhancement of

palynological spectra interpretation. More research on the biological links between sclerotized microfaunal remains and morphologically similar nonpollen palynomorphs in brackish and marine environments is required before precise environmental interpretations can be made from this group of Pleistocene– Holocene palynomorphs. ACKNOWLEDGMENTS The project International Geological Correlation Programme (IGCP) 521 “Black Sea–Mediterranean corridor during the last 30 k.y.: Sea-level change and human adaptation” allowed the co-authors to meet and initiate the idea of this manuscript at the 2007 joint meeting with IGCP 481 “Dating Caspian sea level change.” The senior author acknowledges field-work support from various scientists. Black Sea, Marmara, and Aegean core samples were made available by A.E. Aksu from the archives at Memorial University of Newfoundland (MUN), providing the foundation for this study, and the technical assistance of Helen Gillespie, MUN, for sample processing is gratefully acknowledged. We thank David Horne (Queen Mary University London), Ian Walker (Okanagan University College), Ole Bennike (Geological Survey Denmark and Greenland), and Dirk Verschuren (Ghent University) for help with identifications of animal remains. The Gloeotrichia filament photo (Plate 2, fig. 15) was taken and permission granted for use by Peter A. Siver, at Chrysophytes LLC. Bas van Geel (University of Amsterdam) provided much helpful advice, and we thank J.H. McAndrews (University of Toronto) and J. Matthiessen (Alfred Wegener Institute) for their critical reviews. REFERENCES CITED Aksu, A.E., Yasar, D., and Mudie, P.J., 1995, Paleoclimatic and paleoceanographic conditions leading to development of sapropel layer S1 in the Aegean Sea: Micropaleontological and stable isotope evidence: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 116, p. 71–101, doi: 10.1016/0031-0182(94)00092-M. Aksu, A.E., Abrajano, T., Mudie, P.J., and Yaşar, D., 1999, Organic geochemical and palynological evidence for terrigenous origin of the organic matter in Aegean Sea sapropel S1: Marine Geology, v. 153, p. 303–318, doi: 10.1016/S0025-3227(98)00077-2. Aladin, N.V., and Potts, W.T.W., 1992, Changes in the Aral Sea ecosystem during the period 1960–1990: Hydrobiologia, v. 237, p. 67–79, doi: 10.1007/ BF00016032. Atanassova, J., 2005, Paleoecological setting of the western Black Sea during the past 15000 years: The Holocene, v. 15, no. 4, p. 576–584, doi: 10.1191/0959683605hl832rp. Batten, D.J., 1996, Chapter 26B. Palynofacies and palaeoenvironmental interpretation, in Jansonius, J., and McGregor, D.C., eds., Palynology: Principles and Applications: American Association of Stratigraphic Palynologists Foundation, v. 3, p. 1065–1084. Batten, D.J., and Grenfell, H.R., 1996, Botryococcus, in Jansonius, J., and McGregor, D.C., eds., Palynology: Principles and Applications: American Association of Stratigraphic Palynologists Foundation, v. 1, p. 205–214. Bérard-Therriault, L., Poulin, M., and Bossé, L., 1999, Guide d’Identification du Phytoplankton Marin de l’Estuaire et du Golfe du Saint-Laurent Incuant également Certains Protozaires: Publication Spéciale Canadienne des Sciences Haliétiques et Aquatiques 128, 387 p. Beyens, L., and Meisterfeld, R., 2001, Protozoa: Testate amoebae, in Smol, J.P., Birks, H.J.B, and Last, W.M., eds., Tracking Environmental Change Using

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MANUSCRIPT ACCEPTED BY THE SOCIETY 22 JUNE 2010

Printed in the USA

The Geological Society of America Special Paper 473 2011

Climatic and environmental oscillations in southeastern Ukraine from 30 to 10 ka, inferred from pollen and lithopedology Natalia P. Gerasimenko* Earth Sciences and Geomorphology Department, Taras Shevchenko National University of Kyiv, Glushkova 2, Kyiv, DSP 680, Ukraine

ABSTRACT Pollen and lithopedological data were obtained from Upper Paleolithic sites and Upper Pleistocene loess-soil sequences located between the Sea of Azov and the River Donets, and in the foothills of the Crimean Mountains. During the last Middle Pleniglacial interstadial (Upper Vytachiv soil, 30–27 ka), there existed boreal steppe (south-boreal forest-steppe in Crimea). During the Late Pleniglacial, two main phases of loess accumulation occurred, which were separated by the phase of initial pedogenesis. The loess accumulated under subperiglacial xeric steppe (particularly dry at 15–13 ka), and the incipient soils (Dofinivka unit, 18–15 ka) formed under boreal grassland. During the Late Glacial interstadials, there existed boreal and southboreal forest-steppe with a relatively wet climate (middle Prychernomorsk soil unit, the upper soil 11.8–11.4 ka). During the Younger Dryas, grassland reappeared under a dry and cool climate (10.9–10.5 ka). Paleoclimatic changes demonstrate the same pattern in both studied areas, and they correspond well with Black Sea transgressiveregressive cycles. Regional differences still existed—during all phases, the climate was the mildest in the western foothills of the Crimean Mountains, the coldest in the Donetsk Upland, and the driest near the Sea of Azov.

INTRODUCTION The Upper Pleistocene deposits of southern Ukraine have a distinct cyclical pattern of loess and paleosol alternation, which has been used in the Quaternary stratigraphic framework of Ukraine (Veklich, 1993). According to various studies (Shelkoplyas et al., 1986; Gerasimenko, 1999; Gozhik et al., 2000; Rousseau et al., 2001), the Vytachiv soil unit of the Ukrainian framework corresponds to the Middle Pleniglacial, marine iso-

topic stage 3, whereas the Bug and Prychernomorsk loess units, and the Dofinivka soil unit between them, are correlated with the Late Pleniglacial, MIS 2. In this paper, paleoenvironmental information is based on the results of pedostratigraphic and pollen investigation of the aforementioned late Pleistocene units and the Late Glacial deposits. The sections are located in two regions (Fig. 1): the area between the Sea of Azov and the River Donets, and the foothills of the Crimean Mountains. In the Azov-Donets area, several key

*[email protected]. Gerasimenko, N.P., 2011, Climatic and environmental oscillations in southeastern Ukraine from 30 to 10 ka, inferred from pollen and lithopedology, in Buynevich, I.V., Yanko-Hombach, V., Gilbert, A.S., and Martin, R.E., eds., Geology and Geoarchaeology of the Black Sea Region: Beyond the Flood Hypothesis: Geological Society of America Special Paper 473, p. 117–132, doi: 10.1130/2011.2473(08). For permission to copy, contact [email protected]. © 2011 The Geological Society of America. All rights reserved.

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Pleistocene sites have previously been studied for lithopedology and palynology (Veklich et al., 1973; Artyushenko et al., 1973; Sirenko and Turlo, 1986; Gerasimenko and Pedanyuk, 1991), but, because of inefficient pollen retrieval techniques in previous times, percentage pollen diagrams were plotted only latterly. Data have also been obtained from Final Paleolithic sites (Gerasimenko, 1997a, 1997b; archaeology by Gorelik, 2001), and the

Upper Paleolithic site at Amvrosievka (Gerasimenko, 1997b; archaeology by A. Krotova, 1996). In the Crimean Mountains, after the first investigation of the Zaskal’naya sites (Velichko, 1988; Gubonina, 1985), a multidisciplinary paleoenvironmental study of the Crimean Paleolithic was done (Marks and Chabai, 1998; Chabai and Monigal, 1999; Chabai et al., 2004), and this included investigation

Figure 1. Location map of the sections, discussed in the paper. 1—Kabazi II, 2—Skalisty rock shelter, 3—Buran-Kaya III, 4—Novotroitsk, 5—Amvrosievka, 6—Novoraysk, 7—Rogalik II and XII.

Climatic and environmental oscillations in southeastern Ukraine of small mammals, malacofauna, large mammals (Burke et al., 2004; Markova, 2004; Patou-Mathis, 2004), palynology (Gerasimenko, 1999, 2004, 2007), and absolute dating (Rink et al., 1998; McKinney, 1998, Pettit, 1998). The Late Glacial stratigraphy and paleoenvironments have also been revealed from the Crimean Final Paleolithic sites (Cohen et al., 1996; Yanevich et al., 1996, Gerasimenko, 2004). ENVIRONMENTAL SETTING The Azov-Donets area includes the Pryazov and Donets Uplands (200–350 m above sea level [asl]), coastal plains of the Sea of Azov, and the Donets River alluvial plain (Fig. 1). The uplands are dissected by valleys and gullies to a depth of 100– 200 m. Slopes of the uplands and the plains have a continuous loess cover. The Novoraysk and Novotroitsk sites are sections of low plateaus (180–200 m asl), and Amvrosievka is located in the bottom of a paleogully. Rogalik II and XII are situated on a terrace slope at the lower level of the Donets alluvial plain (70–90 m asl). The Azov-Donets area belongs to the northern subzone of the Ukrainian steppe. From north to south, the average January and July temperatures are–7.5 °C, +22 °C (Rogalik), –7 °C, +21 °C (Novoraysk), –6.5 °C, +22 °C (Novotroitsk), and –6 °C, +22.5 °C (Amvrosievka). The annual precipitation in the Donetsk Upland (Novoraysk and Amvrosievka) is higher than in the Pryazov Upland and the Donets Plain (Novotroitsk and Rogalik): 500–550 and 450–470 mm, respectively. Chernozems (Mollisol), which dominate the soil cover, have a higher humus content and thicker humus layer in the northern part of the area. The steppe coenoses consist of grasses (Stipa, Festuca) and mesophytic herbs (Herbetum mixtum). Arboreal vegetation grows in gullies: oak-ash forest (Quercus robur and Fraxinus excelsior) in the Donets Upland, and scrub in the Pryazov Upland (Malus silvestris, Pyrus communis, Crataegus monogyna, Prunus spinosa, and Ulmus campestris). The highest part of the Donets Upland is regarded as a forest-steppe because oak-maple-lime forest on gray forest soils partially spreads on watersheds here. Pine forest occupies sands on the river terraces. The sections of the Crimean Mountains studied in this paper are located on the slopes of a cuesta, formed in CretaceousPaleogene limestones, at 240–315 m asl. The slope and the sites have a southern exposure. In western Crimea, the Middle Paleolithic open-area site of Kabazi II is situated in the Alma riverbank, 90 m above the water level, and the Late–Final Paleolithic site, Skalisty rock shelter, is 20 m above the water level in the Bodrak River bank. The thick sequence at Kabazi II was accumulated in a sedimentation trap behind a rock resting on the slope bench. In eastern Crimea, the Middle Paleolithic–Neolithic site Buran-Kaya III is a rock shelter in the Burulcha River bank, 10 m above the water level. All sites are located in the low-mountain forest-steppe. The plateau-like tops of the cuesta are covered by meadow-steppe on chernozems. The lower parts of the slopes are occupied by

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arboreal vegetation, the main components of which are oak (Quercus pubescens), hornbeam (Carpinus orientalis), hazelnut (Corylus avellana), and bloody dogwood (Cornus sanguinea). The average January and July temperatures are, respectively, +0.3 °C and +21.5 °C (Kabazi II and Skalisty rock shelter) and –1 °C and +22 °C (Buran-Kaya III). Precipitation in eastern Crimea is much lower (400–450 mm) than in western Crimea (500–550 mm), so forests are thinner, and steppe is drier in the east. The lower forest belt of the Main Mountain Ridge consists of oak and hornbeam, whereas beech and hornbeam grow in the second belt. Pine forests occupy the highest parts of the mountain slopes. METHODS For lithopedological study, humus, CaCO3, and dry salt content, absorbing capacity, and results of bulk chemical and grainsize analyses are interpreted on the basis of the paleopedological approach of Veklich et al. (1979). In order to obtain sufficient pollen counts from subaerial deposits, the following technique was used for sample processing: treatment with 10% HCl, disintegration in a solution of 15% Na4P2O7, treatment with HCl and 10% KOH, cold treatment with HF, and double separation in a heavy solution of CdI2 and KI (specific gravity 2.0 and 2.2). The pollen counts varied between 100 and 500 grains per sample, and pollen was mostly well preserved, particularly in the sections from depressions. Pollen preservation did not depend on carbonate (or gypsum) content in deposits. No palynotypes of Neogene (or older) plants were traced, and in loesses, no pollen of broad-leaved species occurred. This may indicate the absence of pollen redeposition. The transfer functions of vegetation and palynospectra, based on surface samples from different ecosystems, including mountain ones (Grichuk and Zaklinskaya, 1948; Arap, 1976; Dinesman, 1977; Klopotovskaya, 1976; Bezus’ko et al., 1997), were used in the interpretation of pollen data. Pollen percentages were counted from the total sum of microfossils, with the exception of the Buran-Kaya III site (spores were strongly over-represented there, and the percentages of arboreal and nonarboreal taxa were plotted from the arboreal pollen (AP) and nonarboreal pollen (NAP) sum). Pollen of dicotyledonous herbs (with the exception of the Chenopodiaceae and Asteraceae families) is grouped as “Herbetum mixtum”—an indicator of mesophytic steppe type (Grichuk and Zaklinskaya, 1948). In the following sections, the identification of loess-soil units has been primarily made on the basis of lithopedostratigraphy described in the Quaternary framework of Ukraine (Veklich, 1993). The geochronology in loess-soil sections is rather poor and mainly based on 14C data using bones from the archaeological sites. The obtained thermoluminescence (TL) dates, archaeological data, and paleomagnetic age estimations have also been used for the correlation. The geochronological information is summarized in Table 1.

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Gerasimenko TABLE 1. THE GEOCHRONOLOGICAL DATA FOR THE SITES STUDIED

Site

Stratigraphic Archaeological Material Method of Age unit horizon dating (yr B.P.) Buran-Kaya III pc3 4A Bone AMS 10,580 ± 60 Buran-Kaya III pc3 4A Bone AMS 10,920 ± 65 B1 Bone AMS 28,840 ± 460 Buran-Kaya III vt3 B1 Bone AMS 28,520 ± 460 Buran-Kaya III vt3 C Bone AMS 32,350 ± 700 Buran-Kaya III vt2 C Bone AMS 32,200 ± 650 Buran-Kaya III vt2 Buran-Kaya III vt2 C Bone AMS 36,700 ± 1,500 II/1A Tooth ESR 30,000 ± 2,000 Kabazi II vt3b II/1 Bone AMS 31,550 ± 600 Kabazi II vt2 II/2 Bone AMS 35,100 ± 850 Kabazi II vt2 II/4 Bone AMS 32,200 ± 900 Kabazi II vt2 II/5 Bone AMS 33,400 ± 1,000 Kabazi II vt2 II/7AB Tooth ESR 36,000 ± 3,000 Kabazi II vt1c Kabazi II vt1c II/7AB Tooth ESR 38,000 ± 4,000 Novotroitsk vt1b Paleosol TL 57,500 ± 3,000 Novoraysk ud Loess TL 75,000 ± 4,000 III/2 Bone AMS 11,680 ± 110 Rock shelter Skalisty pc2 III/3 Bone AMS 11,750 ± 120 Rock shelter Skalisty pc2 IV Bone AMS 14,570 ± 140 Rock shelter Skalisty df3 Rock shelter Skalisty df3 V Bone AMS 15,550 ± 310 VI Bone AMS 15,030 ± 150 Rock shelter Skalisty df1 Rock shelter Skalisty df1 VII Bone AMS 18,300 ± 220 14 I Bone C 11,400 ± 140 Rogalik XII pc2 I Loess TL 13,500 ± 2,000 Rogalik II pc2 I Paleosol TL 13,000 ± 2,000 Rogalik II pc2 Rogalik II pc2 I Paleosol TL 13,500 ± 1,500 Rogalik II df Paleosol TL 15,500 ± 3,000 Rogalik II df Paleosol TL 17,000 ± 3,000 Peredel’sk bg Loess TL 25,500 ± 2,000 Amvrosiivka df I Bone AMS 18,220 ± 200 Amvrosiivka df I Bone AMS 18,620 ± 220 Amvrosiivka df I Bone AMS 18,700 ± 240 Amvrosiivka df I Bone AMS 18,860 ± 220 Note: AMS—accelerator mass spectrometry; ESR—electron spin resonance; TL—thermoluminescence.

PEDOSTRATIGRAPHY AND PALYNOLOGY Vytachiv (vt) Unit Depending on paleorelief, the Vytachiv unit is represented either by polygenetic “welded” soil or by a pedocomplex of three soils, correlated with the three Middle Pleniglacial interstadials (Gerasimenko, 1999, 2001). The upper paleosol vt3, formed after 30 ka, is examined in this paper (Figs. 2 and 3). This soil is frequently separated from the lower Vytachiv soils by loesslike loam vt2, dated between 36.0 and 31.5 ka, which has low counts of AP (and no pollen of broad-leaved taxa) and high NAP values (Figs. 3A and 3B). The vt2 material is distinguished within the BCca horizon of vt3 soil by much lower contents of R2O3 (Al2O3 and Fe2O3), clay particles, and humus than in the whole Vytachiv pedocomplex (Figs. 2A and 2C). In the DonetsAzov area, vt3 soils are thin, brown-colored, not rich in humus, but they are strongly calcified. The content of humus decreases, whereas contents of CaCO3 and dry salts increase to the south. In the north, water-salt residue in the soil is not considerable and is dominated by Ca-Mg hydrocarbonates (Fig. 2A), whereas in the south, it increases strongly (0.49%–0.60%) and consists of

Reference Chabai et al. (2004) Chabai et al. (2004) Chabai et al. (2004) Chabai et al. (2004) Chabai et al. (2004) Chabai et al. (2004) Chabai et al. (2004) Chabai (2004) Chabai and Monigal (1999) Chabai and Monigal (1999) Chabai and Monigal (1999) Chabai and Monigal (1999) Chabai (2004) Chabai (2004) Gerasimenko and Pedanyuk (1991) Gerasimenko and Pedanyuk (1991) Cohen et al. (1996) Cohen et al. (1996) Cohen et al. (1996) Cohen et al. (1996) Cohen et al. (1996) Cohen et al. (1996) Gorelik (2001) Gerasimenko (1997b) Gerasimenko (1997b) Gerasimenko (1997b) Gerasimenko (1997b) Gerasimenko (1997b) Gerasimenko (1997b) Krotova (1996) Krotova (1996) Krotova (1996) Krotova (1996)

Na-Ca-Mg sulfates. Absorbed Na becomes significant (Fig. 2C), and small gypsum crystals appear. By morphology and properties, vt3 soils are similar to Haplic Kashtanozem in the north and Gypsic Kashtanozem in the south. In the Crimean Mountains, the vt3 unit formed with an admixture of limestone colluvium. It is a brown rendzina (Eutric Leptosol), consisting of a graybrown humic A1 horizon and a light-brown transitional B horizon (Fig. 3), or its pedosediment. Pollen spectra of vt3 soil are dominated by NAP. The AP percentages are higher (23%–44%) in the Donets area and in western Crimea (Figs. 2B and 3A), and lower (4%–13%) in the Pryazov area and in eastern Crimea (Figs. 2D and 3B). Pinus dominates

Figure 2. Sites of the Azov-Donets area (the upper parts of the profiles): (A) Novoraysky quarry, pedolithology; (B) Novoraysky quarry, pollen diagram; (C) Novotroitsk quarry, pedolithology; (D) Novotroitsk quarry, pollen diagram. Legend for Figures 2–7. 1—Chernozem or A1 (humus) soil horizon; 2—Bt soil horizon; 3—Eutric Leptosol (brown rendzina) and its derivatives; 4—Cambisol and its derivatives; 5—Dystric Cambisol; 6—Kashtanozem; 7—Haplic Calcisol; 8—incipient soil; 9—nonsoil loam; 10—loamy pedosediment; 11— sandy pedosediment; 12—loess; 13—loess-like loam; 14—ashy layer.

Climatic and environmental oscillations in southeastern Ukraine

121

A

B

C

D

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B

Figure 3. Pollen diagrams of the Crimean sites: (A) Kabazi II, the upper part of the section (modified from Gerasimenko, 1999); and (B) Buran-Kaya III, the upper part of the section (modified from Gerasimenko, 2004).

A

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Climatic and environmental oscillations in southeastern Ukraine the AP, with the exception of the Buran Kaya III riverine site, in which Alnus prevails. In western Crimea, pollen counts of broadleaved species (Quercus and Carpinus) reach 4%–14%, but no pollen of broad-leaved trees occurs in the Pryazov Plateau, and only a few grains of Quercus or Tilia appear in the other sites. A small amount of Betula pollen is traced in all sections, and single pollen grains of bushes (Corylus, Cornaceae, Elaeagnaceae) are detected in the Donets Upland. Herbetum mixtum dominates the NAP in Crimea, and it shares dominance with Chenopodiaceae (or Asteraceae) in the Donets-Pryazov area. Poaceae and Cyperaceae pollen occur in all sections (Cyperaceae counts are high in Crimea), whereas Artemisia and Chenopodiaceae values are very low (with the exception of Chenopodiaceae in Novoraysk). Ephedra pollen occurs in Crimea. Spore percentages (Bryales and Polypodiaceae) are the highest in the Buran-Kaya III rock shelter and the lowest in the Pryazov Plateau. The top layer of the soil (subunit vt3c) shows a drop in AP pollen percentages and an increase in pollen of xerophytes. Bug (bg) Unit

stone debris, in Crimea. Bug loess has a lower content of humus and clay and, thus, much lower absorbing capacity than the Vytachiv soil. On the contrary, the loess is richer in carbonates and poorer in dry salts, and has higher contents of SiO2 and lower contents of R2O3 than the Vytachiv soil (Figs. 2A and 2C). NAP strongly dominates in the Bug unit (Figs. 2B, 2D, 3A, 3B, and 4). AP incidence falls to 5%–7% (in western Crimea, not lower than 13%–14%) and mostly consists of Pinus. Betula pollen occurs in small numbers, and few grains of Alnus and Juniperus are present in Crimea (Elaeagnaceae and Rhamnaceae at Kabazi II). The NAP is dominated by xerophytes (Chenopodiaceae and Artemisia), and Asteraceae pollen is seen in all sites with the exception of western Crimea. There, pollen of Herbetum mixtum, Poaceae, and Cyperaceae are also significant (Fig. 3A). In the northern Novoraysk site and the Buran-Kaya III rock shelter, Lycopodiaceae and Botrychium boreale spores appear. Dofinivka (df) Unit This unit consists of two to three weakly developed soils, separated by thin loess beds, or of one “welded” soil on the plateau. In the Donets-Pryazov area, the latter (0.5–1.0 m thick) has light-brown color, low humus content, and a very high amount of CaCO3 (Figs. 2A and 2C), frequently in a form of soft carbonate

Archaeological horizon

The Bug unit is a typical loess (1–2 m thick) in the DonetskPryazov area, and a loess-like loam, with a lot of angular lime-

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Figure 4. Pollen occurrence in the Skalisty rock shelter site.

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nodules. At the soil’s lower limit, thin “tongues” indicate former desiccation fissures. The soil does not differ from loess on the basis of its clay contents, but is richer in R2O3, has a higher absorbing capacity (at the expense of absorbed sodium), and a slightly higher amount of dry salts (including sodium sulfates). In southern Ukraine, Dofinivka soil is justly assigned to a brown semidesert type (Veklich et al., 1973), i.e., Haplic Calcisol. In paleogullies, similar soils developed in sandy loams and, thus, are poor in clay and R2O3, but are still calcified (Fig. 5A). In the Crimean Mountains, the Dofinivka unit includes two to three thin beds of brown and gray pedosediments, separated by loess-like loams (Figs. 3B and 4). AP counts in the unit are the lowest (2%–13%) in the Pryazov Upland and eastern Crimea (Figs. 2D and 3B). In western Crimea, they are 25%–55% (Fig. 4), and, in the Donetsk area, increase from the plateau, 17%–31% (Fig. 2B), to paleogullies, 23%–62% (Figs. 2D, 5A, and 6). The lower subunit df1 and the lower brown horizon of the welded soil are richer in AP than the upper subunit df3 and the A1 horizon of the welded soil. Pinus dominates the AP, particularly in the plateau sites. In depressions, the AP is more diverse, especially in subunit df1. In the Donetsk area, Betula pollen is significant, and grains of Betula Nanae et Fruticosae are present in the paleogully (Fig. 6). A few Quercus, Tilia, Corylus, and Picea pollen grains also occur there in subunit df1. In eastern Crimea, pollen of Juniperus, Salix, Betula, and Alnus are represented, and in western Crimea, single grains of diverse broad-leaved taxa appear (only in subunit df1). The NAP composition differs between the sections of plateaus and depressions: in the former, pollen of Asteraceae and Artemisia (or Chenopodiaceae) dominate (Figs. 2B and 2D), and, in the latter, Herbetum mixtum strongly prevails over xerophytes, and Artemisia counts are very low. Spores are not abundant and include Bryales, Polypodiaceae, and, in depressions, Lycopodiaceae.

totype area near the Black Sea, the subunit includes two brown carbonate soils (Veklich, 1993). In the Donetsk area, pc2 soils are leached of CaCO3 (Gerasimenko, 1997a, 1997b). The upper soil is a Dystric Cambisol with dark-gray A1 horizon, brightbrown prismatic B horizon, low CaO content, and signs of clay and R2O3 translocation (Fig. 5A). AP counts are higher in this soil than in the lower one or the thin loess bed between them (Fig. 5B). Pinus and Betula sect. Albae dominate the AP, and Herbetum mixtum prevails in the NAP of the soils. Few Rhamnaceae and Elaeagnaceae pollen grains occur in the lower soil. In the upper one, Alnus, Corylus, Malaceae, and few Picea also appear, and single pollen grains of Quercus have been detected (Gerasimenko, 1997b). In the loess between these soils, the AP is dominated by Betula sect. Nanae et Fruticosae, and pollen counts of xerophytes increase (particularly of Artemisia and Ephedra). In western Crimea, the two beds of pc2 brown pedosediments are also separated by a thin light loam (Fig. 4). Pollen in the pedosediments is small in number, but rich in composition, and includes few grains of Quercus, Tilia, and Corylus in the lower bed, and also Carpinus, Fagus, and Ulmus in the upper one. The thin pc3 loess has pollen spectra dominated by NAP (Figs. 2B and 4). Betula sect. Nanae et Fruticosae prevails in the AP of the Donetsk area (Gerasimenko, 1997b), and no pollen of broad-leaved trees is present in Crimea (only a few Corylus and Rhamnaceae in western Crimea). In the pc3 subunit, pollen counts of xerophytes are much lower than in the pc1 loess and in the loess between the pc2 soils (the Donetsk sites). Pollen percentages of Poaceae and Herbetum mixtum become significant instead, and in Crimea, spores of Polypodiaceae increase in number, and Botrychium boreale almost disappear (Fig. 3B).

Prychernomorsk (pc) Unit

Vytachiv (vt) Unit

The unit consists of two loess subunits (pc1 and pc3), separated by the middle Prychernomorsk (pc2) incipient soils (Veklich, 1993). Because of Holocene erosion, frequently only the thin loess pc1 is traceable. It is somewhat enriched in humus (an impact of Holocene pedogenesis), but its clay content and absorbing capacity are lower than in the soils (Figs. 2A and 2C). The loess is rich in CaCO3 (13%–16%), but not in dry salts. The NAP domination in the pc1 subunit is large or even absolute (Figs. 2B, 2D, and 3B). In the Donets-Pryazov area, pollen of Betula sect. Nanae et Fruticosae is revealed in the loess (Fig. 6). In the Crimean foothills, few grains of Pinus, Juniperus, Betula, and Alnus occur only at the bottom of this subunit. The NAP is dominated by xerophytes in all sites (with the exception of paleogullies, Fig. 6), and an increase in Artemisia pollen is characteristic. Lycopodiaceae and Botrychium boreale spores appear in the Buran-Kaya III rock shelter. The pc2 subunit is dated to the Late Glacial interstadials (Gerasimenko, 1997a, 1997b; Gozhik et al., 2000). In the stra-

The upper soil of the Vytachiv pedocomplex (vt3), which yielded 14C dates 28,800 ± 500 and 28,500 ± 500 yr B.P., electron spin resonance date 30.0 ± 2.0 ka (Figs. 3A and 3B), is correlated with the Denekamp interstadial of Western Europe. The preceding interval, vt2, dated between 31,500 ± 600 and 36,700 ± 1500 yr B.P., had a dry and cold climate: loess-like deposits with low indices of clay weathering formed under grassland both in the eastern Ukraine and in the foothills of the Crimean Mountains. Xerophytic herbs were frequent on the steppe, and only boreal trees occurred in valleys (mainly pine; in the Crimean foothills, also alder and birch). Presently, birch does not grow in Crimea (only sporadically on the highest ridges). Spore plants of boreal climate (Lycopodiaceae and Botrychium boreale) grew in rock shelters of eastern Crimea. Such an environment can be compared with the stadial established for the corresponding time period in Western Europe (Van der Hammen, 1995), in central Ukraine (Gerasimenko, 2001), and in western Ukraine (Bolikhovskaya, 1986, 1995; Haesaerts et al., 2003).

PALEOENVIRONMENTS AND CORRELATION

Figure 5. The Rogalik II site (modified from Gerasimenko, 1997b): (A) pedolithology; and (B) pollen diagram.

B

A

Climatic and environmental oscillations in southeastern Ukraine 125

Archaeological age

Figure 6. Pollen diagram of the lower part of the Amvrosievka site (modified from Gerasimenko, 1997b).

126 Gerasimenko

Climatic and environmental oscillations in southeastern Ukraine During the interstadial vt3, the soils formed under a warmer climate (higher indices of clay weathering). In eastern Ukraine, the spread of Kashtanozems with carbonates high in the profile indicates that the climate still was dry, and its aridity increased to the south, where gypsum and salts intensely accumulated in the soils. Pollen of the Asteraceae family, dominating in the site of Pryazov Upland, includes many palynotypes of Erigeron acer and Helichrysum arenarium (species that can grow on salt soils). Dry grasslands with scattered Elaeagnus bushes occupied the Pryazov area. In the steppe of the Donets area, valleys and gullies included patches of pine forest with some birch, and, in betterprotected places, a few oak, lime, and hazel. In the foothills of the Crimean Mountains, brown rendzina soils (Eutric Leptosols) formed, and broad-leaved species existed in the woodlands, particularly in western Crimea. There, mixed forest, which included oak and hornbeam, expanded much more extensively than in eastern Crimea, where mesophytic steppes strongly dominated, and only a few Quercus pubescens grew in better-protected places. The retreat of boreal trees (particularly of Betula) from the forest and of boreal spore plants (Lycopodiaceae and Botrychium) from rock shelters occurred. Xerophytes almost disappeared from the herbal cover, particularly in western Crimea. This indicates significant climatic humidity. At the end of the interstadial, an extensive advance of steppe and intense suppression of broad-leaved trees marked a transition to the next cold stage. The 14C dates from the top soil of the Vytachiv alluvial suite of the Dnieper terrace are between 28 and 27 ka (Stepanchuk et al., 2004), and these fit with the end of the last Middle Pleniglacial interstadial. The environment in the lower Dnieper valley was cool and similar to the transitional phase detected from the top of soil vt3 at Kabazi II. In the vt3 soil, cultural horizons II/1A and B1-B (Figs. 3A and 3B) include Mousterian artifacts, whereas the underlying vt2 loess, horizon C (Fig. 3B), contains Upper Paleolithic material. It proves a coexistence of Middle and Upper Paleolithic cultures in Crimea and shows that the Crimean Middle Paleolithic lasted considerably longer than in many other areas (Chabai, 2004). Bug (bg) Unit In the study area, this unit yields a range of ages between 27,080 ± 400 and 18,300 ± 220 14C yr B.P. It also yielded a thermoluminescence (TL) date of 25.5 ± 2.0 ka in its lower part (Gerasimenko, 1997b). This enables correlation of the Bug unit with the first half of the Late Pleniglacial. The Bug unit was marked by the strongest loess accumulation of the Pleniglacial. Loess derivatives covered the lower slopes of the Crimean Mountains. The low contents of clay and wide SiO2:R2O2 ratios in the loess are evidence of a drastic decrease in clay weathering, which together with intense accumulation of carbonates and depletion of humus formation indicate a dry and cold climate. This is also confirmed by an increase in input of large angular debris on mountain slopes.

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Pollen data indicate that southern Ukraine was covered by steppe. In the Pryazov Upland, grassland included a high proportion of xerophytic species of Asteraceae. In the Donets Upland and in Crimea, at the beginning of Bug times, the steppe communities included mesophytic herbs. Nevertheless, in Crimea, they were considerably drier than the meadow steppe, which existed at the end of Vytachiv time. During the second half of Bug time, depletion and decrease in the variety of both mesophytic and hygrophytic herbs occurred. Grassland with strong xerophytic components (including Artemisia) was established then, and trees completely abandoned eastern Crimea. This indicates a rising aridity through Bug times—a trend that is also detected in the other areas of Ukraine (Gerasimenko, 2006). Broad-leaved trees disappeared in Bug times, as did Polypodiaceae ferns in western Crimea. Cold-resistant Lycopodiaceae and Botrychium boreale became more abundant in mountain rock shelters and in gullies of the Donetsk area. This proves the existence of a cold (boreal to subperiglacial) climate. Some pine and birch grew in the dissected Donetsk Upland, but the Pryazov Upland was treeless. The Crimean rivers were framed by alder in places. Pine and juniper (in western Crimea also Elaeagnaceae and drought-resistant species of Rhamnaceae) occurred on slopes. The climate was wetter in the foothills of Crimea (and particularly in western Crimea) than in the plains of southern Ukraine. The strong climatic deterioration in Bug times corresponds closely to the harsh climate of the first half of the Late Pleniglacial. Dofinivka (df) Unit The Dofinivka unit has yielded 14C dates in the range 14.0–17.1 ka in its stratotype region (Gozhik et al., 2000). In the studied area, the Dofinivka soils were 14C dated at 14,570 ± 140, 15,550 ± 310, 15,030 ± 150, and 18,300 ± 220 yr B.P. (Cohen et al., 1996), 18,220–18,860 yr B.P. (Krotova, 1996), and TL dated at 15.5 ± 3.0 and 17.0 ± 3.0 ka (Gerasimenko, 1997b). The complexity of the Dofinivka unit shows that it was formed as pulses of incipient soil formation and loess accumulation alternated. In the Azov-Donets area, weak weathering indices and salt-carbonate accumulation in the Dofinivka soils indicate that the climate was strongly continental. The plateaus were covered by dry steppe. Nevertheless, in gullies, the conditions were wetter, and birch-pine woodland spread much wider than during loess accumulation. During df1 times, deep gullies provided a habitat for all mesophytic plants, including arctoboreal Betula sect. Nanae et Fruticosae and a few broad-leaved trees (oak, lime, and hazel). Ecologically different species have been shown to coexist during the late Pleistocene interstadials in other East European sites (Bolikhovskaya, 1986, 1995). During df3 times, the climate became harsher: broad-leaved species disappeared, the role of herbal xerophytes increased, and reduction of woodland occurred (though not as dramatically as during the units of loess accumulation).

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In the foothills of the Crimean Mountains, the climate of Dofinivka times was much wetter than that of Bug times—foreststeppe was established in the west and meadow-steppe in the east of the area. In eastern Crimea, arboreal vegetation was limited to river valleys, where some alder, willow, birch, pine, and juniper grew. Incipient soils, formed under mesophytic steppe, were richer in humus than the soils of the drier Donets-Pryazov area. In western Crimea, during the climatic optimum df1, woodland included an admixture of diverse broad-leaved trees (beech, oak, elm, lime, and maple), which obviously started to spread from refugia. The presence of beech can be seen as evidence that mountain forest belts were in a lower position than at present. In modern surface samples from the Crimean Mountains (Artyushenko and Mishnev, 1978), pollen proportions of broad-leaved trees are much higher than in df1 soils. Thus, even during the optimum, the climate was cooler than present day, and an interstadial environment existed. After df1 times, significant reduction of forest happened and broad-leaved trees retreated. During df3 times, birch-pine woodland included alder and diverse shrubs (Corylus, Sambucus, Rhamnaceae, and Malaceae). During df2 times, pine strongly dominated (only a few birch, juniper, and Rhamnaceae occurred), and thin loess accumulated, indicating a climate similar to that of a stadial, which was colder and drier than the climates of the df3 and df1 soil subunits. Of the last two, df1 times were warmer and wetter, and that was also the case in the Azov-Donets area. A climatic amelioration during the Dofinivka interstadial is reflected in distinctive vegetational changes only in the wetter areas and localities (Crimea and paleogullies). Nevertheless, soil development during Dofinivka times indicates the stability of the sedimentary environment, controlled by a decrease in loess storms, and, thus, by a regional decrease in climatic aridity. The poorly developed Dofinivka soils might be analogues of the set of incipient soils in western Ukraine and Molodova that were 14 C dated between 19.4 and 17.2 ka (Haesaerts et al., 2003). The Dofinivka unit can also be correlated with the Plyussky interstadial of central Russia (16.5–15.0 ka). Cryophytic species of clubmosses and ferns completely disappeared then from the boreal plant communities (Bolikhovskaya, 1995). This was also the case for the rock shelter vegetation in Crimea during Dofinivka times. Prychernomorsk (pc) Unit Loess of the earliest Prychernomorsk subunit pc1 was C dated in the range between 14.0 and 11.9 ka in its stratotype region (Gozhik et al., 2000), and, in the studied area, it fits between 14C 15,500–14,570 and 13,500–11,400 yr B.P. The main phase of loess accumulation occurred during the early Prychernomorsk time. In the Donets-Pryazov area, xerophytic communities of Artemisia, Chenopodiaceae, and Asteraceae predominated, making the landscape similar to a semidesert. Dry ArtemisiaPoaceae steppe occupied eastern Crimea. Sparse arboreal vegetation was restricted to gullies and valleys, and none existed during some time slices. In western Crimea, only drought-resistant 14

pine and juniper formed woodland. All of this indicates an extradry climate. Absence of broad-leaved trees and appearance of arcto-boreal species of Lycopodiaceae and Botrychium boreale in rock shelters are evidence that it was cold even in Crimea. Shrub birches (elements of periglacial vegetation) started to grow abundantly in gullies of the Donetsk Upland. In central Russia, the driest spell of the late Valdai, marked by a strong increase in Artemisia pollen, is related to the Luzhsky stadial, 15–13 ka (Bolikhovskaya, 1995). The pc2 subunit is dated to the Alleröd in its stratotype region near Odessa (Gozhik et al., 2000) and in Crimea (Fig. 4), where it includes the Final Paleolithic cultural layer (Cohen et al., 1996). In many profiles of the Rogalik and Peredel’sk sites (the Donets area), the pc2 pedocomplex consists of two soils, separated by a thin loess (Gerasimenko, 1997b). The upper soil (14C 11,400 ± 140 yr B.P.) corresponds to the Alleröd, the lower soil is marked by a paleomagnetic event, assigned to 13–12 k.y. B.P. (Vigilyanskaya, 1999), the loess includes the Final Paleolithic layer of 10,000 to 9000 B.C. (Gorelik, 2001), and the lower part of the subunit yielded undifferentiated TL dates 13.5 ± 1.5 and 13.5 ± 2.0 ka. Subunit pc2 is correlated with the Late Glacial. During the lower soil formation, birch-pine forest alternated with mesophytic steppe. This marks a sharp environmental change from the dry Late Pleniglacial climate to the interstadial environment, which was cooler but wetter than present day. This interstadial is correlated with the Bölling. In western Ukraine, forest-steppe vegetation of the first Late Glacial interstadial, dated at 12,300 ± 140 to 11,900 ± 230 yr B.P., already included a few broad-leaved species (Ivanova, 1987). The overlying loess bed was formed under typical grassland with considerable share of xerophytes (Artemisia, Ephedra). Arboreal vegetation almost completely disappeared (only a few shrub birches occurred). This stadial, correlated with the Older Dryas, was much colder and drier than the Bölling, but less harsh than the Late Pleniglacial. During the Alleröd interstadial, forest-steppe was reestablished in the Donetsk area. Woody species spread more extensively, and steppe communities included more mesophytic herbs than during the Bölling. Arboreal Rosaceae and Rhamnaceae were abundant. The existence of a “bush steppe” might be indicated, which is typical for the wetter steppe areas. The other evidence of humidity is a small but frequent presence of spruce. Spread of Picea is a typical feature of the Alleröd interstadial in the central part of Eastern Europe (Spiridonova, 1991). Few oaks and hazelnuts appeared in pine-alder-birch forest at the interstadial optimum, after 11,400 ± 140 yr B.P. (Gerasimenko, 1997a). This indicates that the Alleröd was warmer and wetter than the Bölling. Low ridges of the Crimean Mountains were occupied by forest and meadows. At 11,750 ± 120 to 11,680 ± 110 yr B.P. (Cohen et al., 1996), pine, birch, alder, oak, lime, and hazelnut grew in the forest. Later on, its composition became richer and included also hornbeam, beech, and elm. For the whole area, there is evidence that the Alleröd was cooler and wetter than the present time. Boreal plants, including wetloving spruce and alder, penetrated much farther south than

Climatic and environmental oscillations in southeastern Ukraine present day, and the role of broad-leaved trees was lower (particularly in the east). The subunit pc3 corresponds to the Younger Dryas. It was the last interval characterized by the presence of arcto-boreal plants (Betula nana) within depressions of the Donets-Pryazov area. Accumulation of thin loess layers occurred there (Gerasimenko, 1997b). In eastern Crimea, the corresponding layer (14C 10,920 ± 65, 10,580 ± 60 yr B.P.) is marked by the Swiderian culture (Yanevich et al., 1996). Steppe zone occupied the studied area, with the exception of western Crimea. In the north, mesophytic herbs were rather significant, particularly in depressions. Arboreal associations included sparse pine, alder, and arboreal birch. Eastern Crimea was occupied by grassland with high participation of xerophytes, though in river valleys, sparse tree stands from boreal species occurred. In western Crimea, boreal forest-steppe existed (with few Corylus and Frangula). The Younger Dryas was wetter and less harsh than the Older Dryas. CONCLUSIONS In southeastern Ukraine, the time span between 30 and 10 ka generally corresponds to the main period of loess accumulation in the late Pleistocene. Nevertheless, paleopedological and pollen data indicate climatic and environmental oscillations, the alternation of stadials and interstadials, during this time. Its beginning corresponds to the last interstadial of the Vytachiv unit, dated to 30–28 ka, which is the last Middle Pleniglacial (middle Valdai) interstadial. The first half of late Pleniglacial corresponds to Bug times, during which the most loess accumulation occurred (28–18 ka). The second half of the Late Pleniglacial included the two weak Dofinivka interstadials (between 18 and 14.5 ka), separated by a short stadial, and the early Prychernomorsk stadial pc1 (14.5–13 ka). The middle Prychernomorsk time pc2 corresponds to the two Late Glacial interstadials, separated by a short stadial. The late Prychernomorsk stadial pc3 (10.9–10.5 ka) is correlated with the Younger Dryas. The interstadials are represented by paleosols that are less developed than the modern soils. Their pollen is marked by a larger proportion of arboreal taxa and mesophytic herbs than the sediments from stadials. A characteristic feature of interstadial deposits in southeastern Ukraine is the consistent occurrence of a small number of pollen from broad-leaved species in all the sections studied. This might show that broad-leaved trees existed in the gully refugia and increased in number during the interstadials, which had a southern-boreal climate. In the northern Donetsk area, the stadial deposits (mainly loesses) include palynotypes of arcto-boreal shrub (Betula sect. Nanae et Fruticosae), and boreal species of Lycopodiaceae and Botrychium boreale, alongside pollen of xerophytic herbs. This indicates a subperiglacial climate during the loess formation. A combination of dry steppe communities with abundant arcto-boreal plants is typical for the periglacial area of the East European Plain (Bolikhovskaya, 1986, 1995). In the rock shelters of Crimea, an increase in spores of boreal Lycopodiaceae and

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Botrychium boreale, but absence of arboreal cryophytes in the stadial deposits, gives evidence of a northern-boreal climate. Climatic and environmental dynamics in the Donets-Azov area and the foothills of the Crimean Mountains are schematically reconstructed in Figure 7. In Crimea, the interstadials had a southern-boreal climate, with the exception of the interstadial df3, which had a boreal climate. Forest-steppe existed in western Crimea, and mesophytic steppe spread in eastern Crimea. In the Donets-Pryazov area, the interstadials had a boreal climate, though the role of broad-leaved trees evidently increased in refugia. The climate of the interstadial df3 was northern-boreal (broad-leaved trees retreated, and a few shrub birch appeared). During the Late Pleniglacial interstadials, grassland dominated the landscape, and woody plants were restricted to gullies. During the Late Glacial interstadials, forest-steppe occupied the whole area of southeastern Ukraine. The df3 interstadial was the coldest, and the Alleröd was the warmest of the interstadials. The last Vytachiv interstadial vt3 was also rather warm in Crimea, but its transitional phase to Bug times, vt3c, was cool in both areas (boreal and northern-boreal climates). Stadial environments included subperiglacial dry steppe in the Donetsk-Pryazov area and boreal grassland in the foothills of the Crimean Mountains. The latest Late Pleniglacial stadial pc1 was the driest—a semidesert in the Donets-Pryazov area and a dry steppe in eastern Crimea. The stadial pc3, correlated with the Younger Dryas, was wetter in the Donetsk area and warmer in Crimea. In both areas, the beginning of Bug times was wetter than its later half. During the period studied, western Crimea was wetter than eastern Crimea and provided more habitats for arboreal and broad-leaved flora. The ecosystems were generally warmer and wetter in Crimea than in the Donetsk-Pryazov area. This is obviously a reason why the Middle Paleolithic population survived here up to the very end of the Middle Pleniglacial (28 ka), and coexisted with the Upper Paleolithic cultures (Chabai, 2004). This was not a case for the plains area of Ukraine, where the Middle Pleniglacial deposits yield only Upper Paleolithic artifacts. In both the studied areas, paleoclimatic changes demonstrate the same pattern, and they correspond well with transgressiveregressive cycles of the Black Sea. The last Vytachiv interstadial vt3 is correlated with the end of the Tarkhankutian (Surozhian) transgression, which finished at around 31.3 ka (Arslanov et al., 2005), or 30.5 ka BP (Balabanov and Izmailov, 1988). The cold Bug times, when the thickest loess of southern Ukraine was formed, correspond to the Neoeuxinian regression. The latter started at 28.5 ka (Yanko-Hombach et al., 2002). A decrease in aridity during the Dofinivka interstadials can be correlated with the first rise in the Neoeuxinian basin water level. This early Enikalian transgressive phase is related to a period between 17 and 15 ka (Chepalyga, 2002; Murdmaa et al., 2006). The early Prychernomorsk stadial pc1 could correspond then to the deep “post-Enikalian” regression. The middle Prychernomorsk interstadials pc2 occurred at the same time, when the Neoeuxinian transgression became particularly apparent between 13.8 and

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A

B Figure 7. Soil-vegetational dynamics during 30–10 ka: (A) southeastern Ukraine; and (B) foothills of the Crimean Mountains.

Climatic and environmental oscillations in southeastern Ukraine 12.0 ka (Arslanov et al., 2005). The late Prychernomorsk stadial, pc3, correlated with the Younger Dryas, corresponds to the next Neoeuxinian low stand (Chepalyga, 2002; Murdmaa et al., 2006). ACKNOWLEDGMENTS The author is thankful to V. Chabai, A. Yanevich, V. Cohen, A. Krotova (Archaeological Institute of Ukrainian National Academy of Sciences), and to A. Gorelik (Lugansk Pedagogical University) for the chance to participate in the multidisciplinary study of the Paleolithic sites, and also thanks G. Pedanyuk for his cooperation in the investigation of the Donets area. I thank E. Kvavadze and an anonymous reviewer for useful comments and constructive reviews of the manuscript. REFERENCES CITED Arap, R.Ya., 1976, Sporovo-pyltsevye issledovania poverhnostnyh prob pochv rastitel’nyh zon ravninnoy Ukrainy: Kiev, Institut Botaniki Akademii nauk Ukrainy, 25 p. (in Russian). Arslanov, Kh.A., Dolukhanov, P.M., and Gei, N.A., 2005, Climate, Black Sea levels and human settlements in Caucasus littoral 50,000–9,000 years before present, in Yanko-Hombach, V., Buynevich, I., Chivas, A., Gilbert, A, Martin, R., Mudie, P., eds., Extended Abstracts of the First Plenary Meeting and Field Trip of IGCP 521 Project “Black Sea–Mediterranean corridor during the last 30 k.y.: Sea-level change and human adaptation,” (8–15 October 2005, Kadir Has University, Istanbul, Turkey), Ankara, Tübitak, p. 10–11. Artyushenko, A.T., and Mishnev, V.G., 1978, Istoria rastitel’nosti Krymskikh yayl i ikh sklonov v golotsene: Kyiv, Naukova Dumka, 131 p. (in Russian). Artyushenko, A.T., Pashkevich, G.A., Parishkura, S.I., and Kareva, Ye.V., 1973, Paleobotanicheskaya kharakteristika opornykh razrezov chetvertichnykh otlozheniy sredney i yuzhnoy chasti Ukrainy: Kyiv, Naukova Dumka, 96 p. (in Russian). Balabanov, I.P., and Izmailov, Y.A., 1988, Izmenenie urovennogo i gidrokhimicheskogo rezhimov Chernogo i Azovskogo morey za poslednie 20 tysyach let: Vodnye Resursy, v. 6, p. 54–62 (in Russian). Bezus’ko, L.G., Kostylyov, O.V., and Popovich, S.Yu., 1997, Fitotsenologichna interpretatsia palinologichnikh danykh na prykladi Chornomors’kogo biosfernogo zapovidnyka: Ukrainsky Botanichny Zhurnal, v. 54, no. 1, p. 80–86 (in Ukrainian). Bolikhovskaya, N.S., 1986, Paleogeography and stratigraphy of Valdai (Wurm) loesses of the southwestern part of the East-European Plain by palynological data: Annales Universitatis Mariae Curie-Sklodowska (Lublin), Section B, v. XLI, no. 6, p. 111–124. Bolikhovskaya, N.S., 1995, Evolutsia lessovo-pochvennoy formatsii Severnoy Evrazii: Moscow, Moscow University Press, 269 p. (in Russian). Burke, A., Markova, A.K., Mikhailesku, C., and Patou-Mathis, M., 1999, The animal environment of Western Crimea, in Chabai, V., and Monigal, K., eds., The Middle Paleolithic of Western Crimea, vol. 2: Liège, ERAUL, 87, p. 143–151. Chabai, V.P., 2004, Sredny Paleolit Kryma: Simferopol, Shlyah, 323 p. (in Russian). Chabai, V., and Monigal, K., eds., 1999, The Paleolithic of Crimea: II. The Middle Paleolithic of Western Crimea, Volume 2: Etudes et Recherches Archéologiques de l’Université de Liége 87, 260 p. Chabai, V., Monigal, K., and Marks, A., eds., 2004, The Paleolithic of Crimea: III. The Middle Paleolithic and Early Upper Paleolithic of Eastern Crimea: Etudes et Recherches Archéologiques de l’Université de Liége 104, 479 p. Chepalyga, A.L., 2002, Chernoye more, in Velichko, A.A., ed., Dinamika landshaftnykh komponentov i vnutrennikh morskikh baseynov Severnoy Evrazii za poslednie 130 000 let: Moscow, GEOS, p. 170–182 (in Russian). Cohen, V., Gerasimenko, N., Rekovets, L., and Starkin, A., 1996, Chronostratigraphy of rockshelter Skalisty (Crimea): European Prehistory, v. 9, p. 325–358.

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The Geological Society of America Special Paper 473 2011

Late Pleistocene and Holocene paleoenvironments of Crimea: Pollen, soils, geomorphology, and geoarchaeology Carlos E. Cordova* Department of Geography, Oklahoma State University, Stillwater, Oklahoma 74078, USA Natalia P. Gerasimenko Department of Earth Sciences and Geomorphology, Taras Shevchenko University of Kyiv, Glushkova 2, Kyiv, DSP 680, Ukraine Paul H. Lehman Department of Geography, University of Texas at Austin, Austin, Texas 78712, USA Alexander A. Kliukin Cathedra of Geography, Taurida Vernadsky National University, Simferopol, Crimea, Ukraine

ABSTRACT We discuss pollen, soil, geomorphologic, and archaeological records used for reconstructing climatic, biogeographic, and human-environment events in the Crimean Peninsula during the past 130 k.y. Warm and moist conditions conducive to forest growth prevailed during the Eemian Interglacial (marine isotope stage [MIS] 5e). Although sea levels were higher than at present, a review of the stratigraphic and geomorphic data suggests that the peninsula was not detached from the mainland. During the last glacial period (MIS 5d–MIS 2), conditions fluctuated between steppe and tree growth in warmer places during the stadials, and forest-steppe during the interstadials. The Pleistocene–Holocene transition involved forest growth during the Bølling-Allerød interstadials, steppe during the Younger Dryas, and a forest-steppe during the early Holocene. The establishment of the modern Black Sea ca. 7 ka and increasing temperatures led to the formation of the modern vegetation belts, ushering in optimal conditions for the establishment of Neolithic communities. A dry period peaked around 4–3.5 ka, followed by milder conditions that lasted until the colonization of Crimea by Greek farmers during the middle part of the first millennium A.D. Dry conditions at the end of the same millennium led to the abandonment of agriculture and settlement decline. Sea-level oscillations during the late Holocene had an important effect on shoreline configuration, lagoonal systems, coastal wetlands, and human settlements. Data used in this paper were drawn from a number of published papers, mostly in Russian and Ukrainian, as well as records produced by the authors’ research.

*[email protected] Cordova, C.E., Gerasimenko, N.P., Lehman, P.H., and Kliukin, A.A., 2011, Late Pleistocene and Holocene paleoenvironments of Crimea: Pollen, soils, geomorphology, and geoarchaeology, in Buynevich, I.V., Yanko-Hombach, V., Gilbert, A.S., and Martin, R.E., eds., Geology and Geoarchaeology of the Black Sea Region: Beyond the Flood Hypothesis: Geological Society of America Special Paper 473, p. 133–164, doi: 10.1130/2011.2473(09). For permission to copy, contact editing@ geosociety.org. © 2011 The Geological Society of America. All rights reserved.

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INTRODUCTION The Crimean Peninsula occupies an area of ~26,000 km2 bounded on the west and south by the Black Sea, on the east and northeast by the Azov Sea, and on the north by the Sivash (a lagoonal system connected to the Azov Sea) (Fig. 1). At present, Crimea is connected to the Ukrainian mainland by the Perekop, a 25-km-wide strip of land with its highest elevation around 20 m above sea level. The Crimean Peninsula is surrounded by a relatively shallow continental shelf, and as a consequence, significant drops in sea level have exposed extensive areas of land, thereby modifying local climate and vegetation (Cordova, 2007). Conversely, a rise in sea level, such as the one experienced during the Eemian Interglacial (marine isotope stage [MIS] 5e), meant the inundation of low-lying areas, and a possible separation from the mainland (Lazukov et al., 1981; Chabai, 2007). The northern two thirds of the peninsula are composed of flatlands and rolling plains, while the southern extremity is occupied by mountains. The flat, steppe-like region has strong geologic, climatic, and biogeographic similarities with the Ukrainian plains (Berg, 1950; Didukh, 1992). The southern mountainous area shows geological affinities with the Caucasus Mountains and other orogenic belts of Eastern Europe (Мuratov, 1974), as well as biogeographic affinities with the Caucasus, the Balkans, and

Asia Minor (Didukh, 1992). The semi-isolation of the Crimean Mountains from other regional mountainous areas has resulted in a considerable number of endemic species (Biodiversity Support Program, 1999). The lack of glacial activity on its mountains suggests that temperatures in Crimea during the cold stages of the Ice Age were relatively milder than in other regions, creating favorable conditions for tree refugia (Comes and Kadereit, 1998; Gerasimenko, 1999; Morozova and Kozharinov, 2001; Cordova, 2007). Consequently, Crimea has been an important terrain for the study of plant and animal communities and hominin/human paleoecological change during the last interglacial and glacial stages. Several geoarchaeological and paleoecologic projects during the past decade have provided new pieces of information for reconstructing Pleistocene ecosystem change (Chabai et al., 1999, 2004; Gerasimenko, 1999, 2004, 2005a, 2005b, 2007) and Holocene climatic and human-related change (Cordova and Lehman, 2003, 2005, 2006; Cordova, 2007). The objective of this paper is to discuss some of the paleoenvironmental issues investigated through these earlier studies, as well as research developed in the course of the twentieth century. The main points that this paper addresses are: (1) the effects of late Pleistocene climatic change on vegetation, soils, and Paleolithic subsistence in the Crimean Mountains and Piedmont; (2) the climatic and vegetation changes during the Pleistocene-Holocene

45°

33°

34°

35°

Figure 1. The Crimean Peninsula, general topography and geological transects.

36°

Late Pleistocene and Holocene paleoenvironments of Crimea transition; (3) the development of the Holocene landscape in relation to climate fluctuations and human development; and (4) the effects of late Holocene sea-level oscillations on coastal morphology and human settlement.

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Dates of climatic events in the text are often rounded to ka, unless specific dates are referred to, in which case the calibrated yr B.P. or B.C./A.D. is used. GEOMORPHIC AND GEOLOGIC CONTEXTS

APPROACHES AND METHODS The Crimean Peninsula consists of three main physiographic units: the plains, the piedmont, and the mountains (Fig. 1). Areas of rolling plains and low hills break the uniformity of the plains in the Tarkhankutsk and Kerch Peninsulas. The piedmont corresponds to the gentle slope of the Outer Ridge, and the mountains constitute the Inner and Main Ridges. The top of the Main Ridge has a plateau-like summit, known as the Yaila or Yailas (a Turkic term referring to summer pastures), consisting of predominantly calcareous terrain with caves, sinkholes, and uvalas, most of which contain Quaternary sediments. Tectonically, the Crimean Peninsula occupies two megastructures: the Scythian Platform and the Megaanticlinorium of the Crimean Mountains (Мuratov, 1974). The Scythian Platform constitutes the northern part of the peninsula, including the Tarkhankutsk Peninsula and the northern part of the Kerch Peninsula. Prominent structures of the Scythian Platform include the Novoselovsk Rise and the Tarkhankutsk anticline, the Simferopol Rise, the Al’ma Basin, and the Indolo-Kuban Depression. The latter occupies most of the northern and northwestern plains, including the northern part of the Kerch Peninsula. The Megaanticlinorium of the Crimean Mountains, as defined by Мuratov (1974), consists of a series of anticlines and synclines forming a larger upwarp of Mesozoic–Paleogene origin. The Crimean Mountains are the expression of the northern half of the Megaanticlinorium; the southern half is submerged beneath the sea after having slid down along faults (Мuratov, 1974).

The sources of information discussed here include published and unpublished research by the authors of this paper as well as published research papers of other researchers. The localities discussed in detail are those studied by the authors, and these include the mountains and piedmont (for the late Pleistocene vegetationclimate reconstruction), and southwestern Crimea, the Yaila, and the Kerch Peninsula (for the Holocene). Research by other authors was obtained mainly from papers in Russian and Ukrainian, most of which have limited distribution in the west. The topics include marine, alluvial, and loess sequences, as well as research on sealevel changes and the evolution of the Sivash lagoonal system. Lithostratigraphic and chronostratigraphic names are anglicized where possible (e.g., Karangatskii = Karangat). In other cases the Russian, Tatar, or Ukrainian names are retained. In many instances, spelling of names conformed to the most accepted spelling, or the spelling used in the chapters of the compilation by Yanko-Hombach et al. (2007), or from the most common transliteration of the Cyrillic spelling into English. Radiocarbon dates on sections in the figures are presented in radiocarbon yr B.P. (RCYBP). Holocene ages are calibrated to calendar yr B.P. or B.C./A.D., particularly when historical events are discussed. The calibration method used is the Cologne Radiocarbon Calibration and Paleoclimate Package (CalPal, available at http:www.calpal.de). Tables 1, 2, and 3 show the 68% probability range for all radiocarbon dates mentioned in this paper.

TABLE 1. ACCELERATOR MASS SPECTOMETRY DATES FOR THE HERAKLEAN PENINSULA AND THE TARKHANKUTSK PENINSULA Section

14

BBBP-3

Depth (cm) 60–80

Laboratory number Beta-156480

C age (yr B.P.) 3730 ± 40

68% age range (cal yr B.P.) 4014–4144

Bulk carbon

BBBP-2 " " "

50–90 90–100 90–135 590–600

Beta-156478 Beta-127551 Beta-156479 AA48298

5450 ± 40* 8070 ± 40 8550 ± 40 12,028 ± 90

6221–6289 8882–9036 9508–9544 13,793–14,272

Charcoal Bulk carbon Bulk carbon Bulk carbon

AA " " "

25–30 75–90 170–180 310–320

AA48202 Beta-127550 AA48293 AA48294

2907 ± 30 3060 ± 40 5233 ± 40 9499 ± 50

3000–3116 3233–3337 5947–6095 10,707–11,019

Ash Bulk carbon Bulk carbon Bulk carbon

MM2 " "

110–120 245–270 340–350

AA48296 Beta-127552 T-16421A

4637 ± 60 5730 ± 40 8342 ± 70

5315–5458 6479–6602 9259–9438

Bulk carbon Bulk carbon Bulk carbon

BB

185–190

Beta-137094

7000 ± 70

7754–7915

Charcoal

4090 ± 60

4523–4771

Bulk carbon

EF 130–140 Beta-137095 Note: From Cordova and Lehman (2005).

Material dated

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Cordova et al. TABLE 2. ACCELERATOR MASS SPECTOMETRY DATES FOR THE CHËRNAYA RIVER SECTIONS, TARKHANKUTSKAYA BALKA, AND YALTINSKAYA YAILA Section NG2

SB-1

TKB-1 Yaila

14

Depth (cm) 62–74 139–152 230–240 261–274

Laboratory number Beta-127553 Beta-137097 AA48303 Beta-127554

C age (yr B.P.) 1550 ± 40 3530 ± 40 4258 ± 60 5380 ± 40

68% age range (cal yr B.P.) 1402–1504 3744–3868 4720–4863 6088–6259

Material dated

109–123 220–236 292–312 445–461 445–461

AA77565 AA77566 AA77567 AA77568 AA77568

2689 ± 38 4970 ± 64 5082 ± 96 4591 ± 41 5740 ± 50

2771–2839 5646–5839 5718–5922 5150–5430 6486–6648

Bulk carbon Bulk carbon Bulk carbon Bulk carbon Bulk carbon

275

AA48284

4477 ± 39

5028–5252

Charcoal

56–69 90–108 158–170 194–205

AA78717 AA78718 AA78719 AA48275

2034 ± 32 1618 ± 36 3172 ± 39 7823 ± 93

1951–2042 1440–1547 3373–3439 8524–8837

Bulk carbon Humic acids Humic acids Humic acids

Charcoal Bulk carbon Bulk carbon Bulk carbon

TABLE 3. ACCELERATOR MASS SPECTOMETRY AND STANDARD DATES FROM DEPOSITS IN LASPI BAY (LASP 1 SECTION) 14

Unit or Laboratory C age 68% are range layer number (yr B.P.) (cal yr B.P.) 2540 ± 40 2539–2725 Unit 9 -329* Layer A AA48290 2501 ± 35 2514–2725 Layer A AA48290 2781 ± 35 2835–2930 2940 ± 60 3012–3203 Layer A -331* *Reported by Firsov (1972); these are standard radiocarbon dates.

In general terms, the oldest lithological units are exposed along the southern coast, while the younger ones are exposed on the central and northern parts of the peninsula (Fig. 2). The Crimean Megaanticlinorium consists of rocks of Triassic and Jurassic age in its core, and Cretaceous and Tertiary rocks on its exterior (Fig. 2B). The Heraklean Peninsula is formed predominantly by Miocene marine deposits creating a cuesta-like structure, with a veil of terrestrial Pliocene and Pleistocene deposits on top (Fig. 2A). The Kerch Peninsula combines Paleocene– Oligocene rocks in the south with Miocene marine deposits on the north, all of which form a series of synclines and anticlines (Blagovolin, 1962) (Fig. 2C). The Pliocene–Pleistocene geology of Crimea presents several facies, showing first an alternation of loess (or loess-like clays) and paleosols (Blagovolin, 1962; Мuratov, 1974; Veklich and Sirenko, 1976). During the early Pleistocene, the precursors of the modern stream valleys developed, and during the middle and late Pleistocene, a series of fluvial terraces formed within them.

Calibrated calendar age 682 ± 43 BC 653 ± 89 BC 933 ± 47 BC 1158 ± 95 BC

Material dated Timber Hearth Charred wood Charred wood

ing from a higher elevation and orographic precipitation. The southern slopes of the mountains have a relatively warm, subMediterranean climate (Douguedroit and Zimina, 1987; Yena et al., 1996). The Heraklean Peninsula and most of the foothills of the mountains in the southwest possess climatic conditions that are moister than the semiarid steppes and somewhat cooler than the warm southern coast (Cordova and Lehman, 2005; Cordova, 2007). The main weather patterns contributing to the Crimean climatic conditions are the seasonal variation of the Westerly Winds, the Siberian High, and the common cold fronts associated with mid-latitude cyclones. Westerly Winds predominate most of the year, bringing moisture from the adjacent sea, although occasionally, winds from the steppe weaken the effect of the Westerly Winds, particularly north of the mountains. Accordingly, the western and southwestern coasts of Crimea enjoy a relatively wetter climate than the rest of the peninsula.

CLIMATIC AND BIOGEOGRAPHIC CONTEXTS While the northern plains, and the Tarkhankutsk and Kerch Peninsulas, experience a semiarid climate similar to that of the southern Ukrainian and Russian plains, the piedmont and the mountains benefit from a wetter and cooler climate result-

Figure 2. Geological transects and greater lithostratigraphic units. Transect locations are indicated on the map of Figure 1. Data have been assembled from a diverse number of sources (e.g., Blagovolin, 1962; L’vova, 1982; Мuratov, 1974; Podgorodetskiy, 1998; Polkanova and Varuschenko, 1969; Semenenko et al., 1979; Veklich and Sirenko, 1976; Slavin, 1975).

Late Pleistocene and Holocene paleoenvironments of Crimea

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Figure 3. Biogeographic complexes and climate variables. Biogeographic data are based on Biodiversity Support Program (1999). Climatic data were assembled from public records.

138 Cordova et al.

Late Pleistocene and Holocene paleoenvironments of Crimea During the winter, the north and east are directly exposed to blasts of Arctic cold, dry air resulting from the expansion of the Siberian High. In contrast, the southern coast, still under the influence of the Westerly Winds, receives cyclonic waves from the southwestern Black Sea and the Mediterranean Sea (Podgorodetskiy, 1988). During the summer, frontal waves associated with cyclones bring rain to the plains and northern part of the mountains (e.g., Fig. 3, Simferopol climograph). During this time, the mountains create a rain shadow on the southern coast, where precipitation remains relatively low, except for a very short period centered in June (e.g., Fig. 3, Yalta climograph). In essence, the climate of the southern coast, with its the relatively mild winters and dry summers, resembles the typical Mediterranean type of climate (Douguedroit and Zimina, 1987), and this presents one reason why native and introduced Mediterranean vegetation thrives in this region. The northern plains, and the Kerch and Tarkhankutsk Peninsulas have steppe vegetation, grading from meadow steppes in the south to the semiarid and salinized steppes of the northwestern coast, the region around the Sivash, and on the southern coast of the Kerch Peninsula (Fig. 3). Although today, the meadow steppe has largely been replaced by crop fields, the original vegetation according to Podgorodetskiy (1988) had four variants: the Tarkhankutsk steppe, which was dominated by grasses and petrophytic herbs, particularly thyme; the central steppe, which was dominated by grasses and a variety of herbs (i.e., herbetum mixtum); the southern section, which was a meadow steppe with

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shrubs; and the Kerch section of the steppe, which was dominated by patches of grasses and xerophytic herbs, particularly wormwood (Artemisia). The semiarid and salinized steppe is less appropriate for agriculture, and therefore, patches of the original vegetation are still found. The dominant vegetation of the salinized steppe consists of small shrubs such as wormwood and chenopods, as well as a variety of salt-tolerant grasses. Descriptions of vegetation for the rest of the region can be summarized from Didukh (1992) and ordered into a simplified north-south altitudinal transect (Fig. 4). The piedmont and the Inner and Outer Ridges, as well as the Heraklean Peninsula, present mainly forest-steppes, where the dominant tree species are pubescent oak (Quercus pubescens) and eastern hornbeam (Carpinus orientalis). The humid northern slopes of the mountains have dense forests of oak and eastern hornbeam at lower elevations, and hornbeam (Carpinus betulus) and beech (Fagus sylvatica) at higher elevations. The mountain summit, or Yaila, consists of meadow-steppe vegetation with alpine elements and numerous petrophytic herbs and grasses. The forests of the southern slope are dominated by pine (Pinus pallasiana and Pinus kochiana), although stands of beech (Fagus sylvatica) and oaks (mainly Quercus petrae) are not uncommon. The lowest 250 m of the southern slopes consist of pubescent oak and eastern hornbeam woodlands, and sub-Mediterranean elements grouped into communities known as shiblyak and phrygana. The shiblyak is a community of shrubs that includes Juniperus oxycedrus, J. excelsa, Pistacia mutica, Paliurus spina-christi, Arbutus

Figure 4. Simplified vegetation transect across the mountains of Crimea (after Cordova et al., 2001).

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andrachne, Cotinus coggygria, and Ruscus ponticus, with scattered stands of eastern hornbeam, pubescent oak, and Stankevicz’s pine (Pinus pityusae). The phrygana is a community of scrub and herbs of the mint family (Lamiaceae), cistus family (Cistaceae), and Asphodeline spp., among other scrub typical of the Mediterranean region. The distribution of soil types in Crimea is closely associated with the vegetation types. The mountain forest reveals predominantly Rendzina soils on steep limestone and marl slopes, forest brown soils under dense forest on relatively gentle slopes, and meadow soils on the Yaila. The piedmont and plains are characterized by Chernozem (black earth) in the lower part of the piedmont, corresponding to the meadow steppes and open areas of the forest-steppe. The black earth variant of Crimea, which corresponds to the meadow-steppe and open areas of the foreststeppe, is known as the “southern Chernozem of the piedmont” (Krupskiy and Polupan, 1979). The northern steppes of Crimea, which are considerably drier, are characterized by salinized soils such as solonetz and solonchak. In areas around the Sivash, the combination of high water table, low precipitation, and high summer temperatures is the main agent of soil salinization (Podgorodetskiy, 1988; Krupskiy and Polupan, 1979). The southern and southwestern coastal areas tend to have cinnamonic soil, a variation of the terra-rossa identified by Gerasimov (1954) for the sub-Mediterranean parts of the USSR. While cinnamonic soils are common on relatively gentle slopes, stony Rendzina soils predominate on mountainsides. In numerous areas of the western end of the piedmont, however, relict cinnamonic soils exist on gentle slopes. LATE PLEISTOCENE CHRONOSTRATIGRAPHY Marine Deposits A chronology of marine deposits and terraces has been constructed based on the sequences of regressions and transgressions and their correlation with cold and warm stages, respectively. Marine transgression-regression cycles during the late Pleistocene were studied in the Strait of Kerch (Shcherbakov et al., 1977; Fedorov, 1978), the shelf off the southern coast (Shcherbakov et al., 1977), the mouth of the Salgir River (Semenenko et al., 1979), and the western and northwestern shelf (Shcherbakov et al., 1976; Gozhik et al., 1987; Gozhik and Novosel’skiy, 1989). Regressions are commonly marked by alluvial and eolian (mainly loess-like) sediments deposited on the exposed shelf. Transgressions are expressed as marine terraces above the present sea level or by marine sediments backing up stream valleys. Schemes of correlation between MIS (marine isotope stages) and transgressional and regressional events in the northern Black Sea region have also been described by Veklich et al. (1993), Gozhik et al. (2000), and Chepalyga (2007). In the general chronology of sea-level events, the Paleo-Euxinian represents the lowering of sea level accompanied by accumulations of eolian silts related to the Riss Glaciation (Fedorov, 1978), or MIS 6, 8,

and 10 (Chepalyga, 2007). The Karangat transgression has been assigned to the Eemian Interglacial, or MIS 5e (Fedorov, 1978; Veklich et al., 1993; Chepalyga, 2007). The early Neoeuxinian regression is expressed in terrestrial deposits, in limans, and on shelves, as evidence of the coldest stages (MIS 4 and 2), separated by the Surozh transgression (roughly MIS 3 and early MIS 2) (Chepalyga, 1984; Veklich et al., 1993). The Karangat transgression was presumably the highest sea level reached during the late Pleistocene (Fedorov, 1978; Chepalyga, 2007). This was the basis for a suggestion that during the height of the Eemian Interglacial, the Crimean Peninsula became isolated from the mainland by a shallow strait in the Sivash and Perekop areas (Chabai [2007] based on Lazukov et al. [1981]). Reports by Olenkovsky (2000) on prehistoric site survey and in the Sivash area have shown no Middle Paleolithic sites, most of which should correspond to the Karangat and Surozh transgressional stages. On the contrary, several Upper Paleolithic and Mesolithic sites have been found in the area, which suggest that the reduced soil salinity induced by the low sea levels of the Neoeuxinian regression may have made this area more attractive to a diverse assemblage of flora and fauna. The assumption of an island situation during the Karangat transgression is not supported by the stratigraphic record, however. Cores from the Sivash deposits described by Stashchuk et al. (1964) show a series of loess-like and paleosol sequences, but no characteristic marine deposits. Likewise, stratigraphic sections in the Perekop and Solenoye Ozero (Veklich and Sirenko, 1976) show no evidence of marine deposits, but loesspaleosol sequences. Topographic and tectonic factors must be considered in the postulated isolation of Crimea by the Karangat transgression. Accordingly, if the area had been tectonically stable since the time of the transgression, the 7–14 m rise above the present level (Fedorov, 1978; Chepalyga, 1984) would not have inundated the Perekop, where the highest elevations are around 20 m. The Perekop-Sivash area, however, is part of the IndoloKuban Depression, which is constantly sinking, a process that has resulted in the relatively recent inundation of the area and the formation of the Sivash. Therefore, during the Eemian Interglacial, the area should have been higher than it is today, and as a result, the hypothesis regarding the island of Crimea during the Karangat transgression lacks geological support. Sea-level highstands, however, should have created more salinization and a less attractive environment. Loess-Paleosol Sequences The areas with loess and the so-called loess-like deposits occur primarily on the Crimean Plains. Veklich and Sirenko (1976) proposed a chronological classification of Pliocene– Pleistocene silt deposits and paleosols of the western plains near the mouth of the Kacha, Al’ma, and Bulganyak Rivers—the Perekop area. Although the ages of the sequences in the Crimean Plains were not determined through absolute dating methods,

Late Pleistocene and Holocene paleoenvironments of Crimea these deposits and their associated paleosols have been correlated with broad time depositional phases and paleosol formation elsewhere in Ukraine. According to this scheme, the late Pleistocene consists of several pedocomplexes and loess (or loess-like) units that, in turn, are correlated with warm and cold stages, respectively. The Kaydaky soil unit is correlated with the Karangat transgression (Veklich et al., 1993). In the Crimean Plain, this unit is represented by a pedocomplex of two Chernozem soils. The next late Pleistocene soil unit—the Pryluky pedocomplex—includes the upper “chestnut” soil (a soil type of the present dry steppe) and the lower “braunerde.” The Vytachiv soil unit (coeval with the Surozh transgressive phase) is represented by specific “graybrown” soils of the dry steppe, frequently with gypsum and salts. The last late Pleistocene soil unit—the Dofinivka pedocomplex—includes incipient semidesert soils. The loess units that separate these pedocomplexes are generally no thicker than 2 m, with the exception of the Bug loess (coeval with the early Neoeuxinian regression), which reaches 10 m in thickness. Pollen data indicate cold dry steppe during the deposition of all loess units: meadow steppe, steppe, and forest-steppe of temperate climate for the Kaydaky and Pryluky units, and cool steppe climate for the Vytachiv and Dofinivka units (Veklich et al., 1993; Sirenko and Turlo, 1986). Correlation of units established by the Ukrainian stratigraphical framework with their respective marine isotope stages was attempted in later papers (Veklich, 1990; Veklich et al., 1993; Gozhik et al., 2000; Rousseau et al., 2001; Gerasimenko, 1999, 2001, 2006, 2007; Lindner et al., 2002). The scheme of Veklich et al. (1993) does not necessarily match those of other authors (Table 4), but it suits the late Pleistocene pedostratigraphy of the mountains and plains of Crimea. For this reason, this scheme is used in this paper. Alluvial Terraces The chronology of alluvial terraces has been a subject of debate since its preliminary classification by Andrusov (1912), whose work in the Sudak area resulted in a correlation between the drop of the sea in post-Karangat times and the formation of the late Pleistocene terraces. Later on, correlation of the Sudak

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terraces with the river terraces north of the mountains led to a broader scheme of terrace sequences (Мuratov, 1974). This scheme is based on the numbering of terraces from the modern floodplain upward, beginning with terrace I, or the Garden Terrace (Sadovaya Terrassa), of terminal Pleistocene–early Holocene age. Novik (1990) attempted a chronology based on the weathering intensity of clasts in the sediments of the alluvial terraces of the Al’ma River Valley (Table 4). Years earlier, Kliukin and Shchepinskiy (1983) assigned relative ages to terraces II and III in the Al’ma River Valley using diagnostic Paleolithic artifacts. Terrace II has Upper Paleolithic material, while terrace III has late Mousterian (late Middle Paleolithic). Knowing that the boundary between these two lithic industries lies between 40 and 30 ka, the archaeological dates reported by Kliukin and Shchepinskiy (1983) match the dates estimated by Novik (1990) (Table 5). In addition to the lack of absolute dates, a series of problems haunts this chronology of alluvial terraces. These problems involve the local tectonic controls in each valley, which should be considered when cross-correlating terraces among river valleys. Other problems include gradients, catchment size, and elevation in relation to the changing sea levels. One example of the latter is the case of terraces along the Chërnaya River (a special section is presented herein), which presents only one clear terrace above the modern floodplain. It seems that the Chërnaya terraces potentially lie below its recent floodplain, which is barely above the present sea level. VEGETATION AND SOIL DEVELOPMENT Late Pleistocene Vegetation and Soil Development in the Mountains and Piedmont Quaternary stratigraphy and palynology in the mountains and the piedmont areas have been carried out in tandem with archaeological excavations of Paleolithic sites (Chabai et al., 1999, 2004; Gerasimenko, 1999, 2004, 2007). Although many sites are rich in faunal remains, few have pollen records or paleobotanical remains in general. The sites that have provided paleovegetation information are Kiik-Koba (M.N. Klapchuk in Stepanchuk, 2006), Zaskalnaya (Gubonina, 1985; Velichko, 1988), Grot Skalisty (Cohen et al., 1996), Kabazi II (Gerasimenko,

TABLE 4. CORRELATION OF LATE PLEISTOCENE SOIL SEQUENCES AND ISOTOPIC STAGES Unit (index) and equivalent marine Gozhik et al. (2001) Rousseau et al. (2001) Lindner et al. (2002) oxygen isotope stage Prychernomorsk loess (pc) 2 2 2 2 Dofinivka soil (df) 3 2 2 2 Bug loess (bg ) 4 2 2 2 Vytachiv soil (vt) 5a 3 3 3 Uday loess (ud) 5b 4 4 4 5a–c 5c–e 5a–c (or 5) 5 Pryluky soil (pl) Tyasmyn loess (ts) 6 5d (or 6 ) 5d 6 Kaydaky soil (kd) 7 5e (or 7 ) 5e 7 Dnieper loess/glacial unit (dn) 8 6 (or 8 ) 6 8 Potygaylivka soil (pt (zv3)) 9 7 (or 9 ) 7 9 Note: Based on Gerasimenko (2007). Stratigraphic framework is from Veklich et al. (1993)

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Cordova et al. TABLE 5. TERRACE SEQUENCE IN THE AL’MA RIVER AT THE INNER RIDGE WITH WEATHERING-BASED RELATIVE DATES (NOVIK, 1990) AND ARCHAEOLOGICAL DATING (KLIUKIN AND SHCHEPINSKIY, 1983)

Terrace and incision phases Erosion phase High floodplain Erosional phase Terrace I (Garden Terrace) Erosional phase Terrace II Erosional phase Terrace III

Elevation from the channel (m) 0

Estimated age, weathering (ka)

1–2

6.0

1 Present 1

2–4

10.7 –10 .6

1

5–7

31.9

15–25

5 2. 3– 43.6

Erosional phase Terrace IV (Sudak Terrace) 35–45 82–78, 60 Erosional phase 240 , 209 , 151 –1 53 Terrace V (Manddzhil Terrace) Erosional phase Note: Terraces VI to VII are not included in table. MIS—marine oxygen isotope stage. *Broad lithic industry spanning from ca. 130 to 40 ka.

1999, 2005b), and Buran Kaya III (Gerasimenko, 2004). Pollen data in these sites show a conspicuous difference between west (e.g., Kabazi II and Grot Skalistiy) and east (e.g., Zaskalnaya and Buran Kaya) (Fig. 5). This eastward decrease in atmospheric moisture is also apparent in the modern distribution of temperature and precipitation (Fig. 3). Pollen and soil-sediment sequences from Kabazi II and Buran Kaya III provide a source for the reconstruction of vegetation change from the Eemian Interglacial to the end of the Pleistocene (Fig. 6). The Middle Paleolithic open-area site Kabazi II is located on the southern slopes of the Inner Ridge in western Crimea (44°50′N, 34°02′E, and 301 m elevation). The site is located 90 m above the bottom of the Al’ma River valley. The thick sequence of Pleistocene deposits occurs in a sedimentation trap, formed behind a rock slab that fell onto a slope bench. Buran Kaya III (Gerasimenko, 2004, 2007) is located on the northeastern part of the piedmont (45°00′N, 34°25′E; 250 m elevation) on the Burulcha River bank, 10 m above the present water level. It is a rock shelter with a southwestern exposure, and the valley was cut through limestone of the low external ridges of the mountains. Both sites are presently located in the mountain forest-steppe vegetation. The Quaternary Stratigraphical Framework of Ukraine has been used for stratigraphic subdivision in the Crimean Mountains. This framework was elaborated by a research team under the leadership of M.F. Veklich (Veklich et al., 1993). The suggested correlation of Ukrainian units with the chronological scheme of marine isotopic stages and the European paleoclimatic chronology is shown in Table 4 and Figure 6. The detected evolution of vegetation and climate in the western and eastern Crimean foothills, based on pollen and lithopedology, is described next. During the formation of the Kaydaky (kd) unit, a foreststeppe appeared under warm conditions in the western foothills

Estimated age, lithics (ka)

Malinovka I, middle Upper Paleolithic (33–28 ka) Mousterian* (Middle Paleolithic)

MIS

3

4–3

5a–4 7–6

of the Crimean Mountains. Carpinus-Quercus woodland alternated with meadow steppe (mesophytic type). The forest had a mixture of Ulmus (elm) and Tilia (linden) and well-developed shrub undergrowth. This vegetational type is similar to that presently found in the area. The wet lower part of slopes provided an environment for sedges. An abundance of sedges and ironreduction processes (evidenced in the greenish color of the sediments) indicate an excessive ground-moisture supply in the studied locality at Kabazi II. After that, broad-leaved forests declined to the point of total disappearance and were replaced by steppe. As this happened, the pine forest belt expanded, and birch appeared on the northern slope of the mountains. Xerophytic components appeared in the steppe assemblages, and clumps of shrub possibly existed apart from stands of trees. Reduction of broad-leaved species, appearance of birch, and expansion of pine indicate cooling. Presently, birch does not grow in Crimea (it sporadically occurs only in the highest mountain ridges). Later on, the forest-steppe was replaced by a sparse Pinus forest with admixture of broad-leaved trees and rich herb-fern ground cover. The climate was southernboreal, cooler than today. It is also evidenced by the appearance of Fagus, which presently grows in the higher forest belt of the mountains. Humus accumulation was weak under the light forest (A1B horizon of the Mollisol). Mollisols are generally formed under herb vegetation, or under a well-lit forest (with enough light to enable a rich herbal ground cover). Soil humification gradually increased to the end of the interval, along with reduction of the pine population and spread of Poaceae and Cyperaceae. At the end of the Kaydaky unit deposition, the forest-steppe vegetation type became established again. The woodland vegetation was of broad-leaved type (dominated by Quercus). The steppe vegetation was characteristic of the meadow steppe and meadow types, with abundant Cyperaceae and Ranunculaceae.

Figure 5. Late Quaternary stratigraphic sequences and localities mentioned in the text.

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Figure 6. Pedopalynological and geoarchaeological sequence in the northern foothills of the Crimean Mountains: reproduced from Gerasimenko (2007) with modifications. Soils and sediments: (1) meadow soil (Mollisol), (2) Rendzina and brown soil (mollic and eutric leptosol), (3) Luvisol (slope derivative), (4) Cambisol (or pedosediments), (5) loess (light-yellow sediment), (6) light-gray loam, (7) coarse colluvial sediments. Paleoenvironmental dynamics: (8) western Crimea, (9) eastern Crimea.

Late Pleistocene and Holocene paleoenvironments of Crimea Humus accumulated much more strongly than before (A1 horizon of Mollisol). The climate became drier and slightly warmer than in the preceding phase. In sum, during the interval represented by the Kaydaky unit, a temperate stage of an interglacial was replaced by its post-temperate stage. The complete sequence of the Kaydaky unit is not recorded at Kabazi II—first, because the sedimentation started only when the trap was formed behind the fallen rock, and, second, because the upper soil limit is truncated. The first hominin occupation, presumably by Neanderthals, at the Kabazi II site happened during the formation of the A1B horizon of a Mollisol when sparse pine forest spread over the foothills. At the beginning of the Tyasmyn (ts) unit deposition, an intense erosional incision occurred at the site, followed by a thick colluvial accumulation. The colluvial deposits include brown clay beds, pedosediments that are absent in the underlying sedimentary sequence. This represents possible confirmation that the sedimentation break lasted for a long while, and so might correspond to a climatic shift. Phases of strong colluviation correspond with stadial conditions (Antoine et al., 1999; Haesaerts and Mestdagh, 2000). During the formation of the first Pryluky subunit (pl1b1), forest-steppe existed in the foothills, where the soil cover was Luvisols and their slope derivatives. The pollen counts of broadleaved species indicate relatively warm southern-boreal environments for the whole subunit, and the climate was temperate. Broad-leaved forests were dominated by Carpinus orientalis. The steppe associations were mesophytic. Subunit pl1b1 corresponds to a warm interstadial. Its final phase was marked by the appearance of Abies, which could indicate a downhill encroachment of mountain forest belts at the end of the interstadial. The next Pryluky subunit (pl2) was marked by a sharp decline of broad-leaved vegetation. The presence of birch and pine increased, as well as steppe vegetation. The boreal forest-steppe was established, and Artemisia began to appear, although in small numbers. Pedogenic processes were replaced by colluvial accumulation. This subunit corresponds to a stadial in which the climate was cooler and drier than during the preceding interstadial. During the formation of the pl3b2 subunit, humus accumulated under a southern-boreal forest-steppe. Birch and alder coexisted with other broad-leaved species (oak, hornbeam, and elm). Birch expansion indicates that the climate was cooler than today, and the subunit represents an interstadial. The transition to the next cold stage at the end of the interstadial (sharp reduction of arboreal and broad-leaved trees) was related to the final stage of soil formation—subunit pl3c. The characteristic feature of this subunit is the extensive distribution of hygrophytic vegetational components: alder and sedges. A decrease in evaporation, caused by the cooling, could possibly have led to excessive moisture supply at Kabazi II. The Mousterian cultures existed in Crimea during the formation interval of the Pryluky unit. During the deposition of the Uday (ud) unit, pedogenic processes ceased, and light colluvial material accumulated. This was a cold and relatively dry stadial revealing an absence of broad-

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leaved trees and a prevalence of steppe. Boreal trees (pine, birch, and alder) formed groves. Xerophytization of the former meadow steppe is evidenced by the spread of Chenopodiaceae, Artemisia, Ephedra, Plumbaginaceae, and the vanishing of Polypodiaceae ferns in the western foothills of the Crimean mountains. During this same stadial, a xeric steppe occupied eastern Crimea (Velichko, 1988; Gubonina, 1985). At the beginning of the accumulation of the Vytachiv unit (subunit vt1b1), Cambisols and their slope derivatives developed under southern-boreal forest-steppe conditions in western Crimea. Although Pinus dominated the woods, Quercus and Carpinus were common. The appearance of Fagus might indicate that the higher mountain forest belt had moved lower. The climate was cooler than today, and the mesophytication of the herbal cover and reappearance of Polypodiaceae and Lycopodiaceae indicate an increase in precipitation. During the time of accumulation of subunits vt1b1-b2, light colluvial material accumulated instead of soil formation. The mountain foothills were still covered by forest-steppe, but herbal coenoses prevailed, and participation of xerophytes increased considerably (particularly Artemisia and Plumbaginaceae). This, as well as the disappearance of Polypodiaceae and Lycopodiaceae, indicates aridification, whereas the retreat of broad-leaved trees and reappearance of birch give evidence of cooling. Accordingly, subunits vt1b1-b2 corresponds to a stadial. In western Crimea, subunit vt1b2 is marked by a new expansion of broad-leaved trees and mesophytic herbs, and by the development of Cambisols and their derivatives. Pine and birch still prevailed in the forest, and the climate was southern-boreal, characteristic of an interstadial. The domination of woods over steppe, the great herbal diversity, and the spread of ferns indicate an increase in humidity by comparison with the preceding stadial (i.e., MIS 5b). The time of subunit vt1c was marked by a strong expansion of pine forest and a retreat of broad-leaved trees, some of which survived in refugia. Juniperus grew under the light pine stands, the ground cover consisted of mesophytic herbs, and Lycopodiaceae also appeared. The climate was cooler than during the preceding phase and evidently indicated a transition from the interstadial to the next stadial. In the eastern mountain foothills, the forest-steppe environments that existed during the formation of subunit vt1 were much drier than in western Crimea. Meadow steppes (herbetum mixtum and Cyperaceae) strongly dominated over woods, which consisted of boreal trees and a few Tilia. Rock shelters were occupied by diverse spore plants, particularly green mosses. During the formation of subunit vt2, typical grassland was established both in western and eastern Crimea. The woodland was drastically reduced, particularly at the end of the subunit. The western foothills witnessed the first appearance of a steppe vegetation type, and at the end of the interval, the tree population was represented only by alder in the river valleys. In the eastern foothills, trees seem to have disappeared completely. The extensive distribution of Poaceae, xerophytes (Chenopodiaceae, Plumbaginaceae, Artemisia, Ephedra), and Asteraceae occurred.

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Hygrophytes and spore plants grew only in rock shelters and depressions. At the end of the subunit, xerophytes also expanded to the formerly well-wetted localities. Fern and sedge populations strongly declined. The absence of broad-leaved trees and the spread of Lycopodiaceae and Botrychium boreale indicate the boreal climate of this dry stadial, marked by strong colluviation and a complete suppression of pedogenic features. The steppe climate became colder and progressively drier from the beginning to the end of the interval. During the deposition of subunit vt3b, eutric Leptosols were formed, and broad-leaved species reappeared in the woods, particularly in western Crimea. There, a mixture of hornbeam and oak expanded much more extensively than in eastern Crimea, where mesophytic steppes strongly dominated and only a few Quercus pubescens grew in better protected places. The appearance of broad-leaved species and the retreat of boreal trees (particularly Betula) and of boreal spore plants (Lycopodiaceae and Botrychium) indicate an interstadial climate. Xerophytes almost did not exist in the herbal cover during the whole interstadial in western Crimea and at its beginning in eastern Crimea, a condition that indicates relatively high moisture. At the end of this interstadial (subunit vt3c), an extensive advance of steppes and an intense suppression of broad-leaved trees marked the transition to the next cold stage. Thus, during Vytachiv unit times, three interstadials occurred, separated by two stadials. During the first half of the Vytachiv unit, Middle Paleolithic industries occurred in the Crimean Mountains, and starting from the stadial vt2, they coexisted with the Upper Paleolithic occupations in the area (Chabai, 2005). During the formation of this unit, loess-like loams accumulated in the mountain foothills under typical grassland conditions with the substantial participation of xerophytes. Arboreal vegetation was drastically reduced. No broad-leaved species occurred in the valley woods. This indicates the dry boreal climate of a stadial. In western Crimea, Alnus, Betula, and Rhamnus grew in the river valleys, but ferns were completely absent. Drought-resistant and heliophytic plants of the Elaeagnaceae family appeared, indicating the existence of a vast open landscape. In eastern Crimea, at the beginning of the Bug (bg) unit, herbetum mixtum, especially Lamiaceae and Asteraceae, still prevailed in the herb cover, birch and alder trees framed the river, and Pteridae grew in rock shelters. In the second part of the stadial, the area was treeless. Spore plants were dominated by mosses and included boreal elements (club-mosses and grape-ferns). The vegetation assemblages for cold stages (e.g., MIS 4 and MIS 2) show a reduction of broad-leaved species in favor of open steppe, with the prevalence of forest-steppes in some areas (Fig. 6). The presence of boreal trees was evident, but not to the point of dominance. Conversely, the warm stages (e.g., MIS 5e, 5c, 5d, and 3) show the expansion of tree vegetation in the form of forest and forest-steppe communities. The scheme of soil and vegetation change just described refers to the foothills region. No data exist for the higher elevations of the mountains, where environmental conditions and veg-

etation communities during stadials and interstadials can only be inferred. The relict forms of periglacial processes suggest an environment influenced by low temperatures (Podgorodetskiy, 1988), yet the lack of cirque landforms and tills indicates that ice caps and glaciers did not form. Furthermore, some habitation sites are found on the Demerdzhii Yaila at high elevations (often above 800 m), suggesting that conditions may not have reached a full glacial environment. Given the relatively mild conditions of the Crimean Mountains during the coldest periods, ideas for the possible existence of glacial refugia of temperate and boreal trees in Crimea have been put forward based on the presence of haploid types in Fagus (Comes and Kadereit, 1998), stands of relict boreal species in the modern landscape (Morozova and Kozharinov, 2001; Cordova et al., 2001), and modeling and simulation of glacial temperatures (Leroy and Arpe, 2007). The Buran Kaya III and Kabazi II pollen diagrams (Gerasimenko, 2007) show basically an absence of temperate species on the northern slopes. Carpinus is present in small amounts in Kabazi III, but it is more likely to be the more cold tolerant Carpinus betulus. It is probable, however, that glacial refugia for temperate trees were on the warmer southern slope of the mountains. Unfortunately, there are no dated paleobotanical sites on the southern slopes, but the existence of refugia on south-facing slopes of valleys dissecting the northern slopes of the mountains is possible. A study of this aspect, however, requires the use of temperature modeling on slopes using data drawn from the existing pollen record. Faunal assemblages changed in step with vegetation and climate changes. During the coldest stages of the glaciation, the Crimean Mountains had a distinctive mammalian assemblage, distinct from other mountainous regions around the Black Sea (see Markova and Puzachenko, 2006). During the early Weichselian and Denekamp Interstadials, the mountains supported an assemblage represented by woolly mammoth, woolly rhinoceros, wild horse, saiga, red, roe, and giant deer, mountain sheep and goat, cave bear, cave hyena, and lemmings, among others. The Last Glacial Maximum was characterized by European ass, saiga, red and giant deer, northern mole-vole, and steppe and yellow lemmings (Markova and Puzachenko, 2006). During the same period, faunas in the steppes seem to have been dominated by boreal species, strongly represented by woolly mammoth, woolly rhinoceros, reindeer, primitive bison, saiga, arctic fox, cave hyena, and arctic hare. Overall, during these three periods, steppe and forest-steppe fauna dominated, coinciding with the vegetation communities demonstrated by pollen data (Fig. 6). Although boreal and even tundra species existed, they seem to have been of less importance, or perhaps animals associated with the coldest part of the mountains or seasonal immigrants from the boreal areas to the north (Burke et al., 1999). Pleistocene-Holocene Transition After the LGM in the western foothills of the Crimean Mountains, conditions began to favor trees over steppe, as evidenced by

Late Pleistocene and Holocene paleoenvironments of Crimea pollen records from Grot Skalisty (Cohen et al., 1996; Gerasimenko, 2005a). On the eastern foothills of the mountains, however, conditions remained steppic, as evidenced by pollen from zone X (level 5) in Buran Kaya III (Gerasimenko, 2004). Pollen records from the west and east show fluctuations to mesic conditions indicated by a peak in arboreal vegetation consisting of Carpinus (hornbeam), Tilia (lime), and Fagus (beech), suggesting mesic bioclimatic conditions during the Allerød (11,800– 10,900 radiocarbon yr B.P.; 13,900–12,800 cal. yr B.P.). For this interval, pollen records from the Heraklean Peninsula provide evidence for dense forests of Quercus (oak) with a significant presence of Ulmus (elm) and Corylus (hazel), which reflect mild and wet conditions (Cordova and Lehman, 2005). The onset of the Younger Dryas in the Heraklean Peninsula represented a decline in arboreal vegetation (Fig. 7). A transition from a brown forest soil to a meadow Chernozem in geoarchaeological sections parallels pollen evidence for steppic vegetation (Cordova and Lehman, 2005). Reduction of arboreal vegetation occurred also in the western foothills of the mountains, although Rhamnus and Corylus pollen still appeared in Grot Skalistiy (Gerasimenko, 2005a). Farther east, at Buran Kaya III, conditions indicate grass-steppe vegetation. After the Younger Dryas, conditions improved in the foothills of the mountains as the forest-steppe vegetation appeared, even around the eastern side of the foothills at Buran Kaya III (Gerasimenko, 2004). In the Heraklean Peninsula, however, conditions after the Younger Dryas never reverted to forest. Steppe conditions with a few shrubs persisted through the Preboreal and Boreal periods, until roughly 7.4 ka. This development was interpreted by Cordova (2007) as a result of a dry and continental climate maintained by the still low sea level, which continued to influence effective precipitation by an increased continentality. Although the sea level was low during the moist Allerød, lower temperatures, and consequently low evaporation rates, kept conditions mesic. Another explanation might be that the effects of the Westerlies were not strong enough in the southern part of Crimea during this time. It is not until 7.5–6 ka that arboreal pollen appears in larger frequencies in the Heraklean Peninsula (Fig. 7). This development might have been associated with the return of high water levels in the Black Sea, i.e., near modern levels, as well as the overall moister conditions achieved in the region during the Atlantic stage, as evidenced in other areas of southern Ukraine (Kremenetski, 1995; Gerasimenko, 1997). The Pleistocene-Holocene transition encompassed the Mesolithic period, which is strongly represented in practically all regions of the Crimean Peninsula (Bibikov et al., 1994; Olenkovsky, 2000). However, most of the paleoenvironmental and cultural information for the Mesolithic comes from sites in the mountains (e.g., Shan Koba, Fatma Koba, Laspi; Fig. 5), where many lithic and animal bone materials have been recovered (Bibikov et al., 1994; Cohen et al., 1996). Mesolithic sites in the mountains provide macroscopic floral and faunal remains that suggest forest-steppe and forest environments, quite opposite those of the predominantly steppe condi-

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tions of the Upper (Final) Paleolithic (Bibikov et al., 1994). The main aspect in the faunal remains lies in the gradual disappearance of the saiga (Bibikov et al., 1994, p. 167). Unfortunately, beyond the Buran Kaya III and Grot Skalistiy pollen records, little is known of vegetation during the Mesolithic of Crimea. The number of archaeological sites in the Sivash area, as well as the lithic and faunal material recovered in them (Olenkovsky, 2000), suggests conditions more attractive than today’s semiarid climate and salinized soils and water. The early Neolithic cultures of Crimea developed out of the Mesolithic Shan-Koba and Murzak-Koba cultures and appear no later than 8 ka (Cohen, 1996). This emergence coincides with the Boreal period and with cool and dry—but gradually improving— conditions in the southwest. At section BBBP-2, Mesolithic Murzak-Koba lithics are associated with a Chernozem soil and continuing steppic conditions in the Crimean Peninsula (Fig. 8). The later Neolithic cultures, however, had the benefit of relatively warm and moist conditions during the Atlantic period (Fig. 7). The Neolithic meant the beginning of farming and pastoralism, which should be seen as the beginning of a deeper vegetation transformation. The increase of phrygana vegetation in pollen records between 8 and 7 ka (Fig. 7) suggests possible impacts of pastoral activities (Cordova and Lehman, 2003, 2005, 2006). However, the increase in Mediterranean elements in the vegetation assemblages and the possible shift to a dry summer may have played a role in this development. On the other hand, accumulation of sediments in the balkas of the Heraklean Peninsula increased as soils became redder (cinnamonic soils) (Fig. 8). Whether this was the result of climatic shifts or the impact of pastoralism, or a combination of both, is a matter that deserves further research. Middle-Late Holocene in Southwestern Crimea Southwestern Crimea Results obtained through geoarchaeological and palynological studies in southwestern Crimea offer a chronological model that links vegetation-soil development and climate change (Fig. 7). The rise of temperature and moisture that occurred between 7.4 and 6 ka brought about an increase in arboreal vegetation. The apparent reduction of tree cover and increase in nongrass plants suggest a dry phase that peaked between 4.5 and 4 ka. The strong presence of Lamiaceae (mint family) associated with this period suggests that pastoralism may also have been involved in the deterioration of vegetation. Although the pastoral influence is possible, records from the Ukrainian mainland show a drying trend during this period (Kremenetski, 1995; Gerasimenko, 1997). Moist conditions returned by 3.5 ka, and the trend seems to have persisted until around 2 ka (approximately the beginning of the first century A.D.). The forest expansion that took place due to this moist period shows evidence of an interruption: a reduction of tree vegetation coinciding with the establishment of ancient Greek farms (Fig. 7). Pollen records from the floodplain

Figure 7. Pedopalynological composite section for the Heraklean Peninsula and Lower Chërnaya (based on Cordova and Lehman 2005). See Table 6 for designations of Mediterranean, sub-Mediterranean shrub, and Phrygana taxa. PC—post-classic chernozem; MR—meadow rendzina; CC—calcic cinnamonic; BC—brown cinnamonic; Ch—chernozem; BF—alluvial broen forest soil; GS—gray alluvial soil.

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TABLE 6. POLLEN TYPES IN THE GROUPS OF THE POLLEN DIAGRAMS Group Sum of the following pollen types and taxa Broad-leaf trees Quercus, Fagus, Carpinus, Corylus, Pistacia, Tilia, and Ulmus Aquatic and subaquatic Cyperaceae, Typha, Juncus, Sparganium, and Potamogeton Phrygana Cistaceae, Lamiaceae, Asphodeline, Sarcopoterium, Thymelea hirsute Shiblyak Quercus pubescens-type, Pistacia, Jasminum, Paliurus, Cotinus Mediterranean Pistacia, Jasminum, Thymelea hirsuta, Sarcopoterium, Asphodeline, and Cistaceae Sub-Mediterranean Quercus pubescens-type, Cotinus, Cornus, Paliurus, and Verbascum Note: From Cordova and Lehman (2005).

Figure 8. Holocene pedogeoarchaeological sections in ancient farming territories in southwestern Crimea and the Kerch Peninsula. Reproduced from Cordova and Lehman (2005) with modifications and additions. See paleosol types (e.g., MR, PC, CC, MC) on Table 7.

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TABLE 7. PALEOSOL TYPES AND THEIR PEDOGENIC HORIZONS IN THE SECTIONS OF THE HERAKLEAN PENINSULA Soils and paleosols Abbreviation Key characteristics Approximate time span Meadow rendzina MR Organic horizon developed on colluvial material. 3 ka to present 10YR 3/3-4/3-5/2 Medieval to modern Post-Classic Chernozem

PC

Organic horizon with low to moderate carbonate development. 10YR 4/1-5/2

Roman to modern

Calcic cinnamonic

CC

Truncated AB horizon and Bwk (carbonate filaments) horizon with moderate blocky to columnar structure. 10YR 5/1-5/3-5/4

4.2–3.5 ka Bronze Age (on top of soil)

Meadow cinnamonic

MC

A horizon on silty loam with reduction marks. No structure. Dark gray-brown color 10 YR 3/3-4/3

5.5–4.2 ka

Brown cinnamonic

BC

Poorly developed A horizon. Poorly developed Bw (cambic) and Bk (calcic) horizon. Brown color. 7.5YR 4/2-4/4, 10YR 4/3

7.5–5.5 ka

Meadow Chernozem

Ch

Partially truncated dark A horizon on a Bk (carbonate horizon). Massive to weak blocky structure. 10YR 4/1-4/2

10.5–7.5 ka Mesolithic, Murzak Koba lithics (bottom of soil)

Alluvial brown forest

BF

Dark brown A horizon developed on rapidly aggrading silts. Moderate blocky structure. 10YR 4/2-5/3-5/4

12–10.5 ka

Alluvial gray soil

GS

Poorly developed soil on silts and sand. No structure. High gravel content. 2.5 Y 5/4-6/4

12 ka or older

Note: From Cordova and Lehman (2005).

wetlands in the lower Chërnaya Valley show the appearance of cultivated grasses and an increase in weeds and phrygana elements during the establishment of the Greek rural economy during the second half of the first century B.C. (Cordova and Lehman, 2003). Ironically, the impact of Greek farming had little or no effect on rapid sediment accumulation in the ravines of the Heraklean Peninsula and the Chernaya Valley. This contrasts significantly with the period of rapid accumulation presumably associated with the Neolithic. Archaeological records show a decline in farming during the last two centuries B.C. The abandonment of farms occurred during a dry phase that is not represented in the pollen record due to its short duration, but it appears in some geomorphological records discussed later herein. Nonetheless, pollen records show an increase in shrub and tree pollen during the first half of the first millennium A.D., suggesting a decrease in farmed areas (Cordova and Lehman, 2003). During the past 1500 yr, the pollen record gives poor resolution of any changes in vegetation; however, the trend during the last millennium indicates an increase in phrygana and sub-Mediterranean species (Fig. 7). The Yaila and the Mountains The presence of a steppic landscape surrounded by forests on the summit of the Crimean Mountains has prompted biogeographers to propose several hypotheses to explain the treeless nature of the Yaila. In an exhaustive bibliographic compilation

of these hypotheses, Artiushenko and Mishnev (1978) were able to single out at least two main groups of hypotheses. One group states that the Yaila was forested, but gradually lost its cover to pastoral nomads who eliminated the forest over time through burning to create summer pastures. The second group states that the Yaila was always a steppe resulting from local climatic conditions of cold and strong winds, and the hydrological characteristics of its karstic nature. These contrasting views about the origin of the Yaila vegetation motivated Artiushenko and Mishnev (1978) to investigate the case using pollen analysis from soil profiles in each of the Yaila subdivisions: Ai-Petrinskaya Yaila, Yaltinkskaya Yaila, Demerdzhi Yaila, Babugan Yaila, and Dolgorukaya Yaila. Depths of the soil profiles rarely reached 50 cm, and no radiocarbon dates were provided. However, the soil profiles were deemed to have spanned the entire Holocene. Артюшенко and Мишнев’s pollen data from the soil profiles showed that, for the most part, the frequencies of arboreal pollen were minimal, ranging from 10% to 40%. Considering the possibility of wind-transported tree pollen, the amounts of tree pollen, although low, were still an overrepresentation of the tree cover on the Yaila. Artiushenko and Mishnev (1978) eventually favored the climatic reasons for the lack of forests, concluding that the Yaila had been treeless for the entire Holocene. In fact, winds are stronger here than anywhere in Crimea (Artiushenko and Mishnev, 1978), and cryogenic soil processes are common (Вахрушев and Клюкин, 2001), both of which curb

Late Pleistocene and Holocene paleoenvironments of Crimea tree development. Overall, the conclusion points to the existence of an Alpine-type of meadow, as opposed to a deforested area turned into meadow. In reality, small pine trees are scattered in the area, as well as planted trees. Our 2 m core from a depression in the Yaltinskaya Yaila provides a pollen record that further contributes to the paleoecological information of the Yaila (Fig. 9). The bottom date is 7823 ± 93 radiocarbon yr B.P. (cal. 8524–8837 yr B.P.). Throughout the core, the frequency of arboreal pollen is never above 35%, except for two samples: 65% and 57%. A large percentage of the arboreal pollen sum is pine, a characteristic present also in the profiles by Artiushenko and Mishnev (1978), who assumed that most of it was blown in from the areas around the Yaila. Pine pollen is produced in abundance and usually carried very easily by the wind, to the point that large amounts of pine pollen appear even in modern samples of the Heraklean Peninsula, where there are no pine trees (Cordova and Lehman, 2003). In general, the results of our study support the idea of a steppic Yaila at least during the past 8000 yr. Peaks of arboreal vegetation are evident at 50 cm depth. Unfortunately, the dates of this part were reversed. Nonetheless, the peak occurs after 2000 yr B.P., suggesting that this is a climatic event of the subAtlantic that correlates with high arboreal pollen data from the Crimean Peninsula and the Chërnaya Valley during the early centuries of the first millennium A.D. (Cordova and Lehman, 2005). An alternative possibility could be the warm conditions attained during the Medieval Climatic Anomaly (i.e., the Medieval Warm Period), but no other record in the Crimean Peninsula attests to the development of this event. Unfortunately, neither the cores nor other pollen records for the past 2000 yr have the chronological resolution to ascertain events of centennial duration. The Yaltinskaya Yaila pollen diagram, however, shows interesting paleoenvironmental information pointing to possible human influences, namely pastoral groups. The first clue for such potential modifications lies in the accumulation rates and types of sediment. While low sedimentation rates occur between 8000 and ca. 4000 yr B.P., particularly in silt-dominated sediment, afterward this time, sedimentation rates increase, particularly with the inclusion of clay, sand, and small gravel. This change in sedimentation rates (ca. 3500 yr B.P.) coincides with an interval of favorable climate and an increase in herding, which may have led to overgrazing and degradation of meadows on the ridges. In order to explore this hypothesis, we are carrying out a series of tests, including magnetic susceptibility, organic matter, opal phytoliths, burnt grass phytoliths, and microscopic charcoal. THE DYNAMIC ENVIRONMENT OF THE HOLOCENE Soils and Geoarchaeology The geoarchaeological study of the Heraklean Peninsula has shown how soil genesis can impact geoarchaeological and paleoecological interpretation (Figs. 7 and 8). A decline in arboreal pollen (Fig. 7) and a shift from brown forest soil to Chernozem in

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the BBBP sections (Fig. 8) mark the transition from the Allerød to the Younger Dryas. This Chernozem soil appears also in sections MM, AA, and BP. Steppe conditions persisted throughout the Boreal stage; however, the soil and pollen record lack sufficient resolution to reveal changes during the Younger Dryas, Preboreal, and Boreal periods, for which the record still shows semiarid, cool, continental conditions (Cordova and Lehman, 2005; Cordova, 2007). The transition from the dry and cool conditions represented by the Chernozem to a moist-warm marine sub-Mediterranean environment is marked by a reddish brown (cinnamonic) soil dated to the Atlantic Period. The cinnamonic soils in the Black Sea region of the former Soviet Union represent warm and dry summer conditions, as do the Mediterranean terra rossa soils (Gerasimov, 1954). Carbonate accumulations in cinnamonic soils increase toward the top of the sequence, indicating that a decrease in moisture occurred between 4.5 and 4 ka. Moist conditions returned before 3 ka, after which an organic-rich meadow soil and Rendzina appeared, persisting until the present (Cordova, 2007). The studied section at the Tarkhankutskaya Balka in the Kerch Peninsula shows somewhat different soil and geomorphological developments (Fig. 8). Its location in the east of Crimea, away from the moist conditions provided by the Westerly winds, is a clue to understanding climatic changes in the area. Here, modern vegetation and soils suggest persistent steppe conditions and soil salinization. Paleosols typical of brown forest and cinnamonic types are absent in this region. Despite the persistent dry conditions, periods of dry and wet fluctuations are evident in the stratigraphy of the Tarkhankutskaya Balka section. The series of alluvial sediments and soils, which can be associated with several occupations, are divided into four stratigraphic units (Fig. 8). Unit I accumulated at a very slow rate, contrary to what occurred in the Heraklean Peninsula during the past 3000 yr. Floodplain sedimentation increased between the tenth and fourth centuries B.C., evident in the accumulation of unit II. This unit represents a rapid though steady accumulation that had short periods of stability until finally stabilizing in the fourth–third centuries B.C. Unit III is represented by floodplain deposits capped by a dark soil, dated to the third century B.C. Unit IV is an alluvial fill deposited after a period of floodplain erosion. Afterward, a second floodplain erosional phase occurred, shaping the valley into the form we see today. Unit I is characterized by a clayey, organically rich deposit with large amounts of cultural material, altogether suggestive of wet and stable conditions. Unit II represents unstable geomorphic conditions with persistent floods, suggesting drastic wet-dry fluctuations. As evidenced by the regional pollen diagrams, the period of rapid accumulation of init II coincides with a transition from drier to wetter conditions, that is, the transition from the Subboreal to Early Subatlantic (Kremenetski, 1995, 1997, 2003; Gerasimenko, 1997; Cordova and Lehman, 2005). During this period, the site was occupied by a native group from the steppes, which was the first to enter into contact with the Greek colonizers

Figure 9. Yaltinskaya Yaila section with pollen frequencies. AP—arboreal pollen; NAP—nonarboreal pollen.

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Late Pleistocene and Holocene paleoenvironments of Crimea of Panticapaion. The pottery found atop unit I can be assigned to the indigenous Late Bronze Age ceramics. The top of unit II is marked by stable and presumably moist conditions with a strong presence of fourth and third century B.C. Greek pottery. The moment coincides with the expansion of farmland within the Greek Bosporan Kingdom centered at Panticapaion (Fig. 9). The improved climatic conditions of this time are evident in the formation of meadow soils on the Heraklean Peninsula (Fig. 8). It is possible that the influence of these moist conditions may have reached the Kerch Peninsula. Records from the steppes of eastern Ukraine also show a wet climate during this period (Gerasimenko, 2007). Unit III contains water-reworked pottery of the third century B.C. This implies that pottery fragments were transported out of eroded sites and deposited by floodwater at this location, suggesting slope instability and/or perhaps site abandonment. Archaeological records show that at the end of the third century B.C., entire farming areas were abandoned in the city state territories of the Kerch Peninsula (Maslennikov, 1997) and on the western coast of Crimea (Shcheglov, 1978). Floodplain erosion ensued after the deposition of unit III. Pottery of the third century B.C. is embedded in the deposits of unit III, presumably as material removed from abandoned sites. A second incision occurred, eroding the deposits of units I, II, and III, and this event may have occurred soon after the occupation (third century B.C.), suggesting that this may have been a period of dry conditions that favored erosion and incision. An event of erosion and floodplain incision occurred again, followed by the accumulation of unit IV, which contains seventh–twelfth century A.D. pottery. The modern soil formed on this unit, and since then, no incision or accumulation has occurred in Tarkhankutskaya Balka. The geoarchaeological value of soils and sediments lies not only in the information they contain reflecting the conditions of soil genesis and their influence on human occupations, but also the sedimentation/erosion events that may have been related to nonclimatic events. Thus, rapid sedimentation seen in the sediments of the Crimean Peninsula, although related to climatic changes, may also be connected to the beginning of landscape transformation by farmers and herders of the Neolithic and later periods (Cordova and Lehman, 2003, 2006). The Bronze–Iron Age inhabitants of Crimea, or the Tauri of the Greek sources, practiced farming and livestock rearing (Sorochan et al., 2000). The Greek colonists who arrived in Crimea beginning in the seventh century B.C. introduced a specialized agricultural system adapted from their Mediterranean homeland. Beyond the Greek agricultural districts, the indigenous agricultural system continued to be practiced for centuries. It was not until the fourth and third centuries B.C. that agriculture expanded beyond the farming districts of the city states, or chora. Two main areas were eventually developed, one in the east led by the cities of Panticapaion, Nymphaios, and Theodosia, and one in the west led by the city of Chersonesos (Fig. 10). The collapse of agriculture in the western part of

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Crimea seems to have occurred at the end of the third century and beginning of the second century B.C. Archaeologists argue that the main causes of decline were political and military, particularly in view of the strength of the Scythian seminomadic groups of the steppe (Shcheglov, 1978; Sorochan et al., 2000). Lower Chërnaya Valley Another area of focus regarding Holocene geomorphological change is the lower Chërnaya Valley (Fig. 11), which corresponds to the lower reaches of the river basin and the depression that forms Sevastopol Bay. The depression is formed by a fault system that cuts through sediments of Eocene to Miocene age (Fig. 2A). Marine terraces of unknown age border the northern flank of the valley, in the vicinity of Inkerman, properly northwest of the Kalamita Fortress (Fig. 11). At a lower elevation, an alluvial terrace with carbonate sediments borders the basin north and south of the SB-1 location. The degree of carbonate density is suggestive of a Pleistocene age; however, its position within the Pleistocene alluvial terrace scheme (Table 4) is difficult to determine due to a lack of dating. Two cores were taken from the middle and upper part of the floodplain (Fig. 11). Core NG-2 represents mainly overbank flood silts, in some cases associated with soils. A paleosol is clearly visible at the bottom of the sequence, and cumulic soils are visible in other areas. Pollen data from the core reveal a wetland environment, which became stagnant water in levels below the aforementioned clay deposit (Cordova and Lehman, 2003, 2005). The bottom of the sequence was dated at 5380 ± 40 radiocarbon yr B.P. (6088–6259 cal. yr B.P.). Core SB-1 is 460 m long, but at its bottom, a paleosol is dated at about the same age as the bottom paleosol in NG-2 (Fig. 11). Although the two cores seem to cover the same time span, their stratigraphy is somewhat different. The most conspicuous difference is a thick layer of horizontally bedded clay, which is thicker in SB-1, suggesting changes in the preexisting topography during the deposition of this unit. These two cores yield only limited insight into the basin’s Holocene stratigraphy; it is evident that the basin possesses a deeper column of Holocene sediments. Previous engineering cores in the Chërnaya have reported 11.5 m of sediment at Chernomorskoye, just west of SB-1, and 40 m from the mouth of the river. This also shows that sediment accumulation may be associated with sea-level change and tectonics, which is one reason why deeper cores are needed. It is not clear what the effects of late Holocene sea-level change were on the basin, but this is a topic for future study, particularly linking some of the late antique sites in the area. Paleosols, such as the one at the bottom of the sequence, may also suggest buried archaeological sites, which is another geoarchaeological topic to pursue in this area. The relation between sedimentation and late Quaternary sea-level change in the basin becomes complicated if periods of alluvial sedimentation are factored in. Therefore, a project of this

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kind will require extensive coring and digging to study a longer Quaternary sedimentation history.

bottom of the sequence had a paleosol that provided a radiocarbon date of 5597 ± 40 yr B.P. (AA-48302). This finding suggests that terrace I has been partially buried by recent floodplain deposits.

Other Holocene Floodplains Geoarchaeology of Laspi Bay The floodplains of the two major rivers north of the Chërnaya Valley have different morphology, due to their geological and topographic conditions. Unlike the deep graben of the lower Chërnaya (Fig. 2A), these rivers flow through a tilting platform that has caused migrations. The Holocene floodplain sediments of the lower Bel’bek Valley are relatively shallow compared to the lower Chërnaya Valley. In some areas, channel incision has exposed modern sediments that lie barely above the bedrock. The lower Kacha floodplain seems to have a longer history of sedimentation, although no channel cut exposures exist. A trench across the floodplain of the Kacha River next to the Sevastopol-Yevpatoria highway exposed a 2 m sequence of floodplain deposits, however. The top of the sequence had overbank flood deposits associated with twentieth century material. The

Lack of soil formation in geoarchaeological profiles usually means rapid accumulation and landscape instability dominated by slope sediments and wave action. The section exposed on the cliffs of Laspi Bay on the southern coast shows an example of unstable slopes, human occupations, and sealevel changes (Fig. 12). The area was originally studied by Firsov (1972), who obtained the first radiocarbon dates. Although originally the deposits of this section were classified by Firsov into lower and upper terraces, a more detailed study has discovered four Holocene terraces. Because the term terrace in this environment may portray the idea of marine terraces, we use the term “surface” to denote the terrace surfaces originally proposed by L.V. Firsov.

Figure 10. Ancient Greek settlement and areas of farming. Figure is after Kryzhitskii (1997) with modifications by the authors.

155 Figure 11. The lower Chërnaya and sections NG-2 and SB-1.

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Figure 12. Section at Laspi Bay.

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Late Pleistocene and Holocene paleoenvironments of Crimea Surface 1 (S1) corresponds to units 5 and 6, which represent the middle Holocene deposits buried by a late Holocene landslide. The remains of human occupation are associated with the deposition of unit 6 (layer A), which lies on surface 2 (S2). Radiocarbon dates obtained from charred wood and ashes produced ages between the twelfth and the seventh centuries B.C. (Table 3). This dating supports the analysis of pottery associated with it, which is classified as Bronze Age (Firsov, 1972), or pottery that precedes the Greek occupation of southwestern Crimea. A series of three landslide deposits (units 7, 8, and 9) subsequently formed, overlying the occupation layers. A tree log embedded in unit 9 produced a radiocarbon date that placed the landslide deposit in the seventh century B.C. (Table 4). This indicates that the landslides occurred during the occupation or shortly after its abandonment. The surface created by a hiatus in landslide accumulation (unit 9) corresponds to surface 3. Later on, gully erosion dissected this surface. In one of the valleys, alluvial deposits of unit 10 accumulated to create surface 4, which was occupied during and after accumulation during the Middle Ages. One of the buried surfaces of this period corresponds to layer B. After the incision of surface 4 by ravines, accumulation resumed to form the modern alluvial deposits (unit 11). Besides the rapid geomorphic events associated with landslides, stream erosion, and accumulation, the sea-level change is recorded in surface 1, which was originally described by Firsov (1972). Because the occupation surface is barely a meter above the modern sea level, the occupation may have occurred during a regressional phase, most likely the Phanagorian regression. The units below the occupation, namely 5 and 6, consist of a mixture of colluvial and marine deposits, indicating most probably a higher sea-level stand than the preceding transgression. Sea-Level Fluctuations Reconstructions of coastal change have been developed for numerous areas of the Black Sea. The only one that takes data from the Crimean shores is the one published by Shilik (1997), which is supported by evidence collected only from archaeological sites on the coast. Visible shoreline changes in Crimea are not reported for periods prior to 6 ka, since most of the evidence is underwater. In this region, sea-level change information in the early Holocene comes from the boreholes on the northwestern shelf and limans (Shcherbakov et al., 1976; Gozhik et al., 1987), in the Kerch Strait (Fedorov, 1978), and the southern coastal shelf and the Kerch Strait (Shcherbakov et al., 1977). Coastal and alluvial deposits found at various depths show the low levels that preceded the Black Sea transgression prior to 6 ka. The fluctuations after 6 ka are better known through evidence in coastal sediments and erosional benches, where the position of archaeological sites has often been used as a chronological reference. These fluctuations have been arranged into a series of regressions and transgressions, the magnitude of which fluctuates within the range of 8 m. Thus, the general

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scheme includes the Old Black Sea transgression, the Phanagorian regression, the Nymphaean transgression, and the Medieval regression. The main problem is that the constructed curves and phases present discrepancies in the timing and magnitude of the sea-level fluctuations (for example, the curves proposed by Fedorov [1978], Chepalyga [1984], and Shilik [1997]; Fig. 13). Федоров suggests fluctuations of no more than 2 m above and below modern sea level for the past 6 k.y, whereas Shilik shows up to 8 m of sea-level drop during the Phanagorian regression, and according to Chepalyga, the Phanagorian regression dipped to more than 10 m below the present level. Models created for the Kuban and Danube deltas by Kaplin and Selivanov (2004) and Giosan et al. (2006), respectively, show that during the past 5 k.y., the deltas have evolved under the influence of a relatively stable sea with fluctuations within +1.5 and –2.0 m. The curves by Fedorov (1978), Chepalyga (1984), and Shilik (1997) are much simpler than the curve of Balabanov (2007), which shows fluctuations within the major transgressions-regressions. Balabanov established stages (eustatic events on the order of millennia), and phases (fluctuations on the order of centuries) in the history of the Holocene sea-level change. Giosan et al. (2006) attributed these regional discrepancies in sea-level curves to the differences in tectonic activity and hydrostatic rebound following the Holocene infilling of the Black Sea basin. Accordingly, sea-level reconstructions should focus on data obtained locally in order to recognize properly the role of tectonics and local conditions in sea-level correlations. Fedorov’s (1978) and Shilik’s (1997) curves show parallel changes. When compared with Balabanov’s (2007) stage-phaseregression chronology, Shilik’s curve shows a correlation with the regression at the end of the Kalamitan phase (ending around 4000 B.C.) and with the regression following the Old Black Sea transgression. The Phanagorian regression seems to reach its lowest level around 500 B.C. in Balabanov’s scheme and around 100 B.C. in Shilik’s curve. The Nymphaean transgression in Balabanov’s Nymphaean stage is interrupted by a short regression sometime between 600 and 700 A.D., which is not recorded in Shilik’s curve. Balabanov’s scheme does not record the Medieval regression of Shilik’s curve. Three localities with geomorphic, historical, and archaeological landmarks are compared with the scheme of sea-level curves, the Chërnaya floodplain (previously discussed), Laspi Bay, and the Sivash (Fig. 13). The dates from the two cores in the Chernaya floodplain (Fig. 11) show a change from overbank flood silt deposition with cumulic soils to permanent impounding from ca. 5 ka to 3.5 ka, which is also indicated by the increase in aquatic vegetation in the basin (Cordova and Lehman, 2005). This could be the effect of a back-up of seawater from Sevastopol into the floodplain as a result of a rising sea level, probably the Old Black Sea transgression (Fig. 13). Overbank sedimentation with cumulic soils in the floodplain returned and remained until recent centuries. No evidence for a Phanagorian regression and Nymphaean transgression exists in the two cores.

Figure 13. Sea-level curves by Fedorov (1978), Shilik (1997), and periodization by Balabanov (2007) with modifications by the authors.

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Late Pleistocene and Holocene paleoenvironments of Crimea The second kind of evidence comes from the stratigraphic section and archaeological settlement in Laspi Bay (Fig. 12). The chronology developed by Firsov (1972) was later modified on the basis of a recent stratigraphic analyses and more recent radiocarbon dates (Table 2). The third case presented here corresponds to the Sivash, which is more complicated and requires a full analysis of its stratigraphic history. Nonetheless, although these examples seem to agree with the regressions and transgressions, they seem to show no evidence for an 8 m Phanagorian regression as proposed by Shilik (1997). The Sivash Phenomenon The system of lagoons between the Perekop and the Azov Sea presents an interesting aspect of environmental change in which tectonics, sea-level changes, and human history are combined. Although numerous descriptions of sediment and soil stratigraphy have been presented for the region (Stashchuk et al., 1964; Veklich and Sirenko, 1976; Olenkovsky, 2000), no absolute dates for a master chronology exist. The origin and age of the Sivash are two issues that have haunted Quaternary scientists to this day. The Sivash consists of interconnected lagoons of shallow water occupying an area of 2540 km2. Differences in depth, salinity, and mineralogy are notable across the area, and, therefore, Stashchuk et al. (1964) divided the Sivash into western, central, eastern, and southern basins (Fig. 14). Depth varies from less than 0.5 m in the western basin to ~3 m in the southern Sivash. The central and eastern basins have an average maximum of 1 m and 2 m, respectively. The deepest of all (3 m) is reached in the southern part of the southern basin. The Sivash, referred to as the Putrid Lake (Gniloe Ozero, in Russian) is known for salt production, where high water salinity is the result of high evaporation rates. Precipitation in this area is the lowest in Crimea, reaching ~200 mm/yr, where potential evapotranspiration is five times this amount (L’vova, 1982). Inside the Sivash, salinity increases considerably from the mouth eastward and southward; the mean salinity varies from 90‰ to 200‰ in the southern basin, from 80‰ to 270‰ in the central basin, and from16‰ to 260‰ in the western basin (values calculated from percentages reported by Stashchuk et al., 1964, p. 12–15). Just for comparison, the salinity of the Azov Sea is 10‰–11‰ (Stashchuk et al., 1964), which is considerably lower than the mean salinity of the world ocean (35‰) due to the influx of freshwater from rivers. Sediment cores from the Sivash bed show first a series of loess-like silt and paleosols, covered by a relatively thin sequence of lagoonal sediments (Stashchuk et al., 1964). This sequence indicates that the Sivash is a relatively recent phenomenon, particularly in the western and central basins, and probably of Holocene origin (Podgorodetskiy, 1988). Archaeological survey of the Sivash area shows that the Upper Paleolithic (40–16 ka) and the Mesolithic (14–8 ka) are the most common settlements in the area (Olenkovsky, 2000). The middle and late Holocene sites are very rare,

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suggesting that interest in the Sivash by human groups became minimal after the Mesolithic period (ending around 7.5 ka). Using historical maps and descriptions since Hellenistic and Roman times, Stashchuk et al. (1964) identified changes in the landscape for the past 2000 yr, particularly for the Arabatskaya Strelka and the western basin. The Arabatskaya Strelka does not appear on maps and historical accounts until the sixteenth century. Before then, the present location of the Arabatskaya Strelka was occupied by two islands in the early first millennium A.D., and then by a chain of islands during the late first millennium and early second millennium A.D. (Stashchuk et al., 1964). This development suggests that the Arabatskaya Strelka formed as an original accumulation of sand along a spit that eventually connected the small islands. Strabo’s descriptions and later Claudius Ptolemy’s map refer to Lake Biki, which occupied the present area of the western Sivash (Stashchuk et al., 1964, p. 160–161). Additionally, a series of shallow lakes, probably similar to the ones existing today south of the Perekop, were often reported in the area of the central basin. The eastern and southern basins were part of the Sea of Azov, until they were separated from it by the Arabatskaya Strelka. Historical observations of the coastal configuration of the Sivash region can be seen in the context of sea-level changes in the late Holocene (Fig. 13). The observations of Strabo and Pliny the Elder were made at the end of the Phanagorian regression, perhaps when large shallow areas had been exposed as islands. In the historical data collected by Stashchuk et al. (1964), there is no report of a sand spit for most of the first millennium A.D.— roughly coinciding with the Nymphaean transgression. The first report of the Arabatskaya Strelka occurs at the end of the seventeenth and more clearly during the eighteenth century, that is, at the beginning of the last regression. The other possible effects of sea-level change on the formation of the Arabatskaya Strelka are shown by Zenkovich (1958), who proved that the accumulation of so great a size and volume of sand and shells in the bar of Arabatskaya Strelka (estimated at more than 300 million m3) could have occurred only over a long period, with the area continuously being exposed and drowned. This scenario is based on the existence of a submerged sand bar parallel to the Arabatskaya Strelka (Fig. 14A). This bar is ~2 km wide, and its crest is 4 m below the sea surface (Zenkovich, 1958, p. 171). Sand along the bar is continuously being accumulated by currents on its west side, consequently producing a westward migration of the entire bar (Zenkovich, 1958, p. 171). Based on the development of the offshore bar, Зенкович came to the conclusion that the Arabatskaya Strelka migrated westward until it reached an elevation high enough to be exposed and become attached to the mainland to the south, thus creating the eastern and southern lagoonal basins of the Sivash. This scenario does not explain the flooding of the central and western basins, however. Therefore, a more complicated scenario involving tectonics, sea level, coastal accumulation processes, and even fluvial input of sediment should be considered.

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Figure 14. Morphological changes in the Sivash region over the past two millennia according to Stashchuk et al. (1964).

Late Pleistocene and Holocene paleoenvironments of Crimea PATTERNS OF ENVIRONMENTAL CHANGE: A CONCLUSION This paper pursued the objective of identifying the most important paleoenvironmental events in the Quaternary of Crimea over the last 130 k.y. using data from research published in previous decades as well as recent work by the authors in order to explore several aspects of late Pleistocene environmental change in the Crimean Peninsula. These include changes in vegetation, soils, sedimentation, and coastlines in relation to climatic changes, and the relations among humans, vegetation, and climate, and the need for higher resolution in the chronology of terrestrial and marine deposits. Coastal changes related to fluctuations of the Black Sea and the Sea of Azov have been important in shaping landform patterns and sediment deposition directly by changing the base level of streams to expose large tracts of shelf or drowning modern valleys and inlets. Evidence for sea-level stands appears in the marine terraces, particularly the Karangat terraces, and in the filling of marine sediments of river valleys, as in the case of the Lower Salgir. In turn, conditions inland were warmer and wetter, in part because these changes occurred during interglacials and interstadials, and in part because the marine sources increased as the peninsula became surrounded by water. The question whether Crimea became an island during the Eemian Interglacial, or MIS 5e, as proposed by Chabai (2007), is still debatable, particularly given the lack of tangible stratigraphic evidence in the Perekop and Sivash area. Vegetation and soil development correlate with the sequence of stadials and interstadials during the Pleistocene. Pollen data point to the predominance of steppe vegetation during cold periods, and forest-steppe and forest during warm periods, a relation that is confirmed by faunal assemblages. The PleistoceneHolocene transition is characterized by the return of forest and forest-steppe communities, which was the dominant vegetation during the Eemian Interglacial. This transition was delayed by oscillations between cold and warm periods, however, and the attainment of modern forest communities in the Heraklean Peninsula and the foothills of the mountains occurred after 7.4 ka, at which time the Black Sea reached levels close to those of today. A dry period centered between 4.5 and 4 ka is apparent in most regional pollen records. The return of humid conditions after 3.5 ka seems to have coincided with the spread of Bronze and Iron Age communities and finally the Greek colonization between 3 and 2.5 ka. Records show little resolution for the past millennium. The influence of humans on the geomorphic and ecological processes of the peninsula can be assumed by coupling geoarchaeological and paleoecological records. During the middle and late Holocene, the Heraklean and Kerch Peninsulas underwent cycles of erosion. Although these cycles might have been climatically driven, they occurred at a time when farming and pastoralism arose in Crimea. In the pollen and soil records of the Heraklean Peninsula, the beginning of the Neolithic is represented by an increase in phrygana scrub. Although an increase

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of phrygana is often related to pastoral activities, it may have resulted from climatic conditions. The effects of the establishment of Greek farming are more clearly seen in the records, particularly in the pollen data. Ironically, there is no evidence for increased erosion. It is not until the abandonment of farms that erosion seems to occur, particularly in the Tarkhankutskaya Balka in the Kerch Peninsula. Although coring and studies of sections have been undertaken for decades, a tighter chronology is needed. In particular, these studies are needed for marine deposits to determine whether they correspond to the Karangat or Surozh transgressions. Dated loess sequences in the plains of Crimea would complement paleoenvironmental data already obtained in the Pleistocene deposits of the mountains. The terraces of the large river valleys are another source of relevant information, particularly through dating and correlation with sea-level changes. Holocene deposits such as those of the Sivash, bays, and lakes are still in need of further study and better understanding. In many cases, data from pollen, diatoms, and other proxy records can also add interesting information to the climatic context in areas where the most recent cultures of the peninsula have evolved. ACKNOWLEDGMENTS This research was funded in part by a grant from the Packard Humanities Institute to the Institute of Classical Archaeology (ICA) of the University of Texas at Austin and small grants by the College of Arts and Sciences at Oklahoma State University. Field research was carried out during summer seasons of years 1998–2001. We thank the numerous members of the staff of the Natural Preserve of Tauric-Chersonesos (Sevastopol), who helped this project in many ways and for providing us with a base for field research. We are thankful to many members of the Faculty of Geography at Taurida Vernadsky National University in Simferopol and the Nikitsky Botanical Gardens in Yalta, and the Institute of Geography of the National Ukrainian Academy of Sciences in Kyiv. REFERENCES CITED (Russian and Ukrainian references are immediately followed by their translations.) Andrusov, N.I., 1912, Terraces of the Sudak Area, Zapiski Kievskogo Oshschestvo Estestvopyt, v. 22, no. 1, no paging (in Russian). Aндpycoв, H.И., 1912, Teppacы oкpecтнocтeй Cyдaкa, Зaпиcки Kиeвcкoгo Oбщecтвa Ecтecтвoиcпыт, v. 22, issue 2. Antoine, P., Rousseau, D.-D., Lautridou, J.-P., and Hatté, C., 1999, Last interglacial-glacial climatic cycle in loess-palaeosol successions of northwestern France: Boreas, v. 28, p. 551–563, doi: 10.1080/030094899422046. Artiushenko, А.Т., and Mishnev, В.Г., 1978, History of the vegetation on the Yailas and surrounding slopes during the Holocene: Kiev, Naukova Dumka, 140 p (in Russian). Apтюшeнкo, and Mишнeв, 1978, Иcтopия pacтитeльнocть Kpымcкиx Яйл и пpияйлинcкиx cклoнoв в Гoлoцeнe: Kiev, Hayкoвa Дyмкa, 140 p. Balabanov, I.P., 2007, Holocene sea-level changes of the Black Sea, in YankoHombach, V., Gilbert, A.S., Panin, N., and Dolukhanov, P.M., eds., The Black Sea Flood Question: Changes in Coastline, Climate, and Human Settlement: Dordrecht, Springer, p. 711–730.

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Veklich, M.F., and Sirenko, N.A. 1976, The Pliocene and Pleistocene of the left bank of the Dnieper and the plains of Crimea: Kiev, Naukova Dumka, 186 p. (in Russian). Beклич, M.Ф., and Cиpeнкo, H.A., 1976, Плиoцeн и Плeйcтoцeн лeвoбepeжья нижнeгo Днeпpa и paвниннoгo Kpымa: Kiev, Hayкoвa Дyмкa, 186 p. Veklich, M.F., Sirenko, N.A., Matviishina, Zh.N., et al. 1993, Stratigraphic scheme of the Pleistocene deposits of Crimea: Kiev: Goskomitet geologii Ukrainy, 40 p. (in Russian). Beклич M.Ф., Cиpeнкo, H.A., Maтвиишинa, Ж.H., et al., 1993, Cтpaтигpaфичecкaя cxeмa плeйcтoцeнoвыx oтлoжeний Yкpaины: Kiev, Гocкoмитeт гeoлoгии Yкpaины, 40 p. Velichko, A.A., 1988, Geoecology of the Mousterian in East Europe and adjacent areas, in Otte, M., ed., L’Homme de Neandertal, L’Environment, Volume 2: Etudes et Recherches Archéologiques de l’Université de Liège, v. 29, p. 181–206. Yena, V.G., Tverdokhlebov, I.T., and Shantyr’, S.P., 1996, The Southern Coast of Crimea: Simferopol, Biznes-Inform, 304 p. (in Russian). Eнa, B.Г., Tвepдoxлeбoв, И.T., and Шaнтыpь, C.П., 1996, Южный Бepeг Kpымa: Simferopol, Бизнec-Инфopм, 304 p. Yanko-Hombach, V., Gilbert, A.S., Panin, N., and Dolukhanov, P.M., eds., 2007, The Black Sea Flood Question: Changes in Coastline, Climate, and Human Settlement: Dordrecht, Springer, 971 p. Zenkovich, V.P., 1958, The coasts of the Black and Azov Seas: Moscow, State Publishing of Geographic Literature, 373 p. (in Russian). Зeнкoвич, B.П., 1958, Бepeгa Чëpнoгo и Aзoвcкoгo Mopeй: Mocквa, Гocyдapcтвeннoe Издaтeльcтвo Гeoгpaфичecкoй Литepaтypы, 373 p.

MANUSCRIPT ACCEPTED BY THE SOCIETY 22 JUNE 2010

Printed in the USA

The Geological Society of America Special Paper 473 2011

Bedforms, coastal-trapped waves, and scour process observations from the continental shelf of the northern Black Sea A. Trembanis* S. Nebel A. Skarke Department of Geological Sciences, University of Delaware, 109 Penny Hall, Newark, Delaware 19716, USA D.F. Coleman R.D. Ballard Graduate School of Oceanography, University of Rhode Island, South Ferry Road, Narragansett, Rhode Island 02882, USA A. Yankovsky Marine Science Program and Department of Geological Sciences, University of South Carolina, Columbia, South Carolina 29208, USA I.V. Buynevich* Geology & Geophysics Department, Woods Hole Oceanographic Institution, MS 22, Woods Hole, Massachusetts 02543, USA S. Voronov Department of Underwater Heritage, Institute for Archaeology, Academy of Sciences of Ukraine, Kyiv, Ukraine

ABSTRACT The Black Sea basin presents an ideal laboratory for investigations of morphodynamic interplay between response (morphology) and force (processes) associated with shelf sedimentation. Recent studies along the perimeter of the basin have documented the existence of a complex, heterogeneous seafloor varyingly composed of sand, gravel, silt, and clay. Side-scan sonar data are utilized to establish the spatial patterns of bedform types in the area. In addition, a benthic tripod, configured with an acoustic Doppler current profiler, a rotary fanbeam sonar, and a conductivity-temperature sensor was deployed to record seabed dynamics in response to changing forcing conditions. Together, the tripod and side-scan survey data sets provide a complementary basis for deciphering the processes responsible for the observed seafloor morphology. The side-scan sonar data allows for the determination of spatial patterns of bedform length and orientation. In total, 2376 individual large sand wave bedforms were

*E-mail, Trembanis—[email protected]; present address, Buynevich—Department of Earth and Environmental Science, Temple University, 1901 N. 13th Street, Philadelphia, Pennsylvania 19122, USA; [email protected]. Trembanis, A., Nebel, S., Skarke, A., Coleman, D.F., Ballard, R.D., Yankovsky, A., Buynevich, I.V., and Voronov, S., 2011, Bedforms, coastal-trapped waves, and scour process observations from the continental shelf of the northern Black Sea, in Buynevich, I.V., Yanko-Hombach, V., Gilbert, A.S., and Martin, R.E., eds., Geology and Geoarchaeology of the Black Sea Region: Beyond the Flood Hypothesis: Geological Society of America Special Paper 473, p. 165–178, doi: 10.1130/2011.2473(10). For permission to copy, contact [email protected]. ©2011 The Geological Society of America. All rights reserved.

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Trembanis et al. digitized in geographic information systems with mean and modal wavelengths of 72.8 and 15.7 m respectively. The correlation of near-inertial waves (velocity amplitude 12–20 cm/s and period 12–16 h) and bedform geometry suggest that the extensive sand-wave patches imaged across the shelf are affected by active modern processes and may themselves be modern features or perhaps relict features that remain active presently. Progressive vector diagrams of the nearbed mean current flow indicate a component of cross-shelf directed flow, suggesting an enhanced potential for artifact preservation via cross-shelf advection of anoxic bottom waters by the near-inertial flows measured in this study.

BACKGROUND The Black Sea provides an ideal natural laboratory for testing the role of shelf transport processes on bedform and artifact interaction (Özsoy and Ünlüata, 1997; Neretin et al., 2001; Coleman and Ballard 2004). This unique setting presents opportunities to test concepts of artifact-related scour and transport associated with complex sorted bedform features (Murray and Thieler, 2004; Trembanis et al., 2004; Green et al., 2004) and shipwreck artifacts (McNinch et al., 2006). The widespread occurrence and previous documentation of large-scale bedforms on the continental shelf (Ryan et al., 1997; Coleman and Ballard 2004; Lericolais et al., 2006) raises the question of whether these bedforms are (1) strictly modern; (2) ancient relicts; or (3) palimpsest features (i.e., relict but reworked) features. The objective of this study is to examine how the hydrodynamics (mean currents) of the shelf interact with the seafloor morphology over spatial scales ranging from meters to kilometers in a morphodynamic context similar to that used in other shelf studies (e.g., Trembanis et al., 2004; McNinch et al., 2006). In addition to analyzing the distribution of bedforms (size and orientation), a secondary goal of this study is to relate the shelf hydrodynamic processes (internal waves) and seafloor morphology (bedforms) as important causative factors in shipwreck site and artifact preservation sensu (McNinch et al., 2006) with application to recent marine archaeological expeditions in the region (Coleman and Ballard, 2004). It has been hypothesized (Ryan et al., 1997; Coleman and Ballard, 2004) that interfacial internal waves assist in transporting anoxic waters onto and across the shelf providing a mechanism for enhanced artifact preservation above the normal oxycline. At the center of these hypotheses is the interplay between hydrodynamics and bed roughness. Previous observations (Ryan et al., 1997; Coleman and Ballard, 2004; Lericolais et al., 2006) suggest that there are strong analogs between the Black Sea shelf settings and the ubiquitous shelf sand body features originally termed “Rippled Scour Depressions” (Cacchione et al., 1984) that have been termed “Sorted Bedforms” in recent years (Murray and Thieler 2004; Trembanis et al., 2004). Of particular parallelism was the finding in New Zealand (Hume et al., 2003; Trembanis et al., 2004) that anoxic organic material in the sediment may have played a stabilizing role in controlling the lateral stability of the bedform features. It is possible that such deposits exist in the vicinity of

the wreck sites in the Black Sea and perhaps these shear resistant deposits play a similarly important role in artifact and site preservation, whereby reduced organic layers might cap and help preserve the artifacts. SHELF MORPHODYNAMICS Field observations and theoretical refinements by numerous investigators over the past several decades have significantly advanced our understanding of shelf sediment transport processes (Thieler et al., 1995; Wright, 1995; Grant and Madsen, 1986). The continental shelf is an important transition region for physical, biological, and geological processes—one that forms a critical link between the nearshore and the deep-sea basin. The shelf is a morphodynamic system influenced by coupled physical, geological, chemical, and biological processes (Wright, 1995). Processes and phenomena of the shelf exhibit strong spatial and temporal variability, making this a complex four-dimensional region of study (Wright, 1995). Within this relatively shallow setting, frictional forces are important in connecting hydrodynamics to the behavior of seabed forms of varying scale (Grant and Madsen, 1986). In a strongly bidirectional manner, the bottom boundary layer structure depends heavily on the morphology of the seabed that in turn is shaped by gradients in the hydrodynamics (Wright, 1995). Furthermore, numerous studies of shelf settings around the world (Trembanis et al., 2004; Schwab et al., 2000; Drake, 1999; Wright et al., 1999; Riggs et al., 1998; Thieler et al., 1995; Cacchione and Drake, 1990) have documented that complexity is more the norm than the exception. RIPPLES AND BEDFORMS Ripples and other large bedforms on the shelf (e.g., sandwaves and subaqueous dunes) are important sources of seabed roughness to waves and currents (Ardhuin et al., 2002; Grant and Madsen, 1986) and play a key role in the nature and magnitude of sediment resuspension (Traykovski et al., 1999; Li et al., 1996; Cacchione and Drake, 1990). Several field and laboratory studies have been conducted in attempts to develop empirical formulae between ripple geometry (e.g., height, length, steepness) and flow conditions (Wiberg and Harris, 1994; Wikramanayake, 1993; Clifton and Dingler, 1984; Grant and Madsen, 1986; Nielsen, 1981; Miller and Komar, 1980). Under the typically irregular

Bedforms, coastal-trapped waves, and scour process observations flow conditions encountered in the field, the ability of these models to accurately predict observed ripple geometry has been shown to be rather poor (Trembanis and Traykovski, 2005; Trembanis et al., 2004; Doucette, and O’Donoghue 2002; Traykovski et al., 1999; Li et al., 1996; and Osborne and Vincent, 1993). In part, the reason for the poor agreement between field data and model estimates is that ripple geometries encountered in the field are often partially relict products of forcings from past events and not solely products of instantaneous hydrodynamic conditions. Both Traykovski et al. (1999) and Li and Amos (1999) observed significant hysteresis in ripple development on the shelf. In addition to non-equilibrium evolution, spatially varying grain size is another important issue affecting ripple dynamics (Green et al., 2004; Trembanis et al., 2004). Grain size and ripple dimensions on the shelf often exhibit large variations over spatial domains both greater than 1 km (e.g., Green et al., 2004; Hume et al., 2000; Black and Oldman, 1999; Barnhardt et al., 1998; and Field and Roy, 1984) and less than 1 km (e.g., Trembanis et al., 2004; Ardhuin et al., 2002; Thieler et al., 1995; Hunter et al., 1988; and Schwab and Molnia, 1987), often in correlation with sorted bedforms (previously termed “rippled scour depressions”) that have wavelengths much longer than the ripples themselves (Green et al., 2004; Traykovski and Goff, 2003). According to Holland et al. (2003), heterogeneous patches of contrasting sediment grain size are frequently encountered along shelf environments around the world. SCOUR Scour is the morphodynamic response of the seabed as a result of the presence of an object or structure that disturbs the fluid flow (Soulsby, 1998). Scour is important for a variety of marine situations including bridge piers, dock pilings, breakwaters, oil platforms, offshore pipelines, marine artifacts, heterogeneous seabed bedforms, and naval mines (Whitehouse, 1998). The presence of an object on the seabed produces local flow acceleration due to continuity and thus drives a flux of local sediment and concomitant bed adjustments (Whitehouse, 1998). Another manifestation of scour is an increase in bed shear stress and turbulence as structured vortices are generated and released from around the object (Trembanis et al., 2007; McNinch et al., 2006). Scour can be classified both in terms of spatial extent and hydrodynamics. Three spatial classes of scour are defined as: “local scour” which is in the immediate vicinity of the object (on the order of meters), “global scour” composed of wide depressions around large or multiple objects (on the order of 10s of meters), and “overall seabed movement” associated with large scale (100s–1000s of meters) patterns of erosion, deposition, and bedform movement (Whitehouse, 1998). In terms of hydrodynamic intensity, scour is classified as either clear-water, when the ambient flow (bed shear stress) is below threshold velocity, or live-bed, when ambient flow is above threshold velocity and the entire bed is active. In the former, the amplification of flow

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about the object induces transport locally but elsewhere the bed is immobile. In the latter case, sediment is being transported by flow everywhere, but especially near the object, where turbulence and bed shear stress are enhanced (Trembanis et al., 2007; McNinch et al., 2006). Once an object is exposed on the seabed, scour is initiated around the lateral ends of the object because of converging accelerated flow. This convergence leads to progressive erosion of the sediment from under the ends of the object, forming an expanding scour pit and shrinking the support pedestal. The object then will settle into its scour pit in a series of rocking and rolling motions until it is no longer protruding above the ambient seabed or until flow conditions subside and backfilling (deposition) ensues (see McNinch et al., 2006, and figures therein). In non-steady flows, periods of excavation (scour) will normally be interrupted by episodes of backfilling (deposition) (Trembanis et al., 2007; Richardson and Traykovski, 2002; Fredsoe, 1978). A possible sequence for scour around a free settling horizontal object, such as those examined in this study, is illustrated in McNinch et al. (2006; their fig. 4). HYDRODYNAMICS AND BEDFORMS Like no other large body of water in the world, the Black Sea (Fig. 1) has an upper oxygenated layer and a lower anoxic layer. The interface between the layers reaches down to 180 m depth along the coastal margins, and 500 m near the Bosporus (Özsoy and Ünlüata, 1997; Neretin et al., 2001). This interface appears to be unstable and at varying times, probably during severe weather conditions, creates internal waves that break upon the Black Sea’s continental shelf along the oxic/anoxic boundary. The surface layer in the Black Sea averages 18‰ salinity and 22 °C, and is highly oxygenated. The surface circulation features a cyclonic rim current about the entire basin. Two cyclonic gyres occur within the outer rim current, as well as eddies and intermittent convection to intermediate depths by surface cooling. The transition from the surface to the denser, trapped deep layer is marked by an oxycline with steep gradients in salinity, temperature, and chemical content. The deep-water averages 22‰ salinity and 8 °C and is completely anoxic while being rich in hydrogen sulfide (H2S) and ammonium (Özsoy and Ünlüata, 1997). The density contrast between these layers is large enough to support the propagation of internal waves. Internal waves with periods of six minutes and more, and with bottom velocities on the order of 30 cm/s, fast enough to exhume and carry suspended silt and fine sand (Hjulström, 1935; Prothero and Schwab, 2004), have been observed on the southern coast of Crimea (Filonov, 2000). When the crest of such a wave meets the continental shelf like that of the Danube delta in the northwest, the anoxic, H2S-rich water runs up and down the slope in a manner precisely analogous to the swash and backswash of a surface wave on a beach—whether in linear or turbulent fashion—and almost certainly has a large effect on the benthic environment. The anoxia and high H2S content of the wave water would kill most or all organisms it

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contacts. In certain areas bottom topography may contribute to a rapid enough backswash to cause sediment entrainment, erosion, and the development of bedforms (McPhee-Shaw, 2006). Internal waves have two important geoarchaeological effects. The first is the creation of a mixed layer between 85 and 185 m that periodically leads to anoxic water conditions. These anoxic water conditions are of interest to archaeologists since they lead to the long-term preservation of ancient wooden shipwrecks (Coleman and Ballard, 2004). The second is the creation of bedforms characterized by large sand waves that lay in a depth zone of 85–185 m. These bedforms are of interest because: • the existence of these shelf-sorted bedform features is poorly documented and not well understood; • the hydrodynamics responsible for these bedforms may help transport anoxic water across the shelf; and • the introduction of anoxic water onto the continental shelf by internal waves may preserve ancient wooden shipwrecks at far shallower depths than previously thought. FIELD SITE AND RESEARCH METHODS The field site was located off the southwest coast of the Crimean Peninsula of the Northern Black Sea (Fig. 1). In

2006, a geoacoustical survey was conducted of the southwestern Crimean shelf and slope (Fig. 2) along a suspected deepsea trade route between the Bosporus and the Crimea. A number of side-scan sonar targets were identified that were later inspected with a remotely operated vehicle. Several of the targets turned out to be modern shipwrecks and aircraft (mostly Russian from the Crimean War, World War I, and World War II eras), but one of the targets, located ~23 km off the coast, was a pile of ancient ceramic jars. Based on the typology of similar-looking jars from the Chersonesos site (Ryzhov and Sedikova, 1999) and elsewhere around the Black Sea region (e.g., Hayes, 1992), these jar types are estimated to date between the ninth and eleventh centuries C.E., placing the ship in the early Medieval Period. At a depth of ~150 m, this shipwreck lies above the normal anoxic interface but within the mixed layer of temporally varying dissolved oxygen content. It has thus escaped the ravages of marine borers and appears to be in an unusually good state of preservation. In general, this mixed layer is prevalent throughout the entire Black Sea basin along the shelf break between depths of ~80–180 m (Ballard et al., 2001). The side-scan sonar data collected in 2006 forms the basis of the large-scale bedform mapping presented in this paper (Figs. 2 and 3).

Scale: 300 km

Scale:

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Figure 1. Study-site location map (star) off the SW coast of the Crimean peninsula at a local depth of ~135 m. Inset illustrating the side-scan sonar surveys in the vicinity of the bottom mooring associated with the shipwreck site known as Chersonesos A (star).

Bedforms, coastal-trapped waves, and scour process observations SENSOR CONFIGURATION AND SAMPLING REGIME In addition to the side-scan sonar surveys conducted in 2006 (Figs. 2 and 3), a set of bottom-mounted instruments (Fig. 4) were deployed in 2007 measuring currents, temperature, salinity, and seabed geometry. The bottom-mount (Fig. 4A) was located at a depth of 135 m and included an upwardlooking acoustic Doppler current profiler (ADCP), conductivitytemperature (CT) logger, and rotary fanbeam sonar. The CT sensor provides a time-series point measurement of the ambient salinity and temperature used to determine changes to the density of the surrounding water and the passage of thermocline or pycnocline oscillations (Fig. 6). The ADCP gathers vertical profiles of three-dimensional hydrodynamic flow (Fig. 7) recorded in earth-coordinates (i.e., east-west, north-south, up-down) based on an internal compass. The rotary fanbeam sonar (Fig. 4B) obtains a high-resolution planview image of the seabed surrounding the bottom mount to a range of 9 m thus providing a

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time-lapse picture (Fig. 5) of the small-scale bedform geometry and dynamics in for comparison to the hydrodynamic measurements. Table 1 summarizes the sampling scheme settings for each of the bottom-mount instruments. On 15 August 2007, the ADCP/CT/Sonar bottom mount was deployed in the vicinity of the Chersonesos A wrecksite at a local depth of 135 m and was recovered after ~38 h. Upon recovery, data was downloaded and archived. The data analysis results are presented below. RESULTS AND DISCUSSION Salinity and Temperature Record Figure 6 illustrates the recorded time series of temperature (blue line) in °C and salinity (green line) in ‰ at a height of 0.75 m above the bed. The Sea-Bird SBE-37SM sensor recorded every 23 seconds, the fastest sampling scheme that would cover the expected deployment duration. The sharp drop in temperature

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C 4 km Figure 2. Map plot illustrating several hundred of the more than 2300 bedforms digitized in geographic information systems. The crest line of each visible bedform was digitized from the side-scan sonar survey data. Insets are as follows: (A) Cape Sarych and the surrounding vicinity; (B) Digitized bedform crests in a 10 km box surrounding the Chersonesos A wreck site; (C) Close-up portion of the survey site showing digitized bedforms (blue lines) over the side-scan sonar record. Star illustrates location of the acoustic Doppler current profiler/conductivity-temperature sensor/Rotary sonar bottom mooring.

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and rise in salinity at the beginning of the deployment represents the measurements as the instrument mount was descending through the water column during deployment. A reverse signal is seen at the end of the deployment when the bottom mount was recovered from the seabed. During the interval between descent and recovery, the CT sensor recorded essentially static values for salinity and temperature. The mean temperature was 8.3 °C with a standard deviation of 0.03 °C. Mean salinity was 20.6‰ with a standard deviation of 0.08‰. The slight changes in the values during the deployment and the sharp changes at the beginning and the end confirm that the sensor was in fact working and that the nearly flat line plots of temperature and salinity were in fact real. We initially expected sharp fluctuations in temperature and salinity coinciding with the interface of the pycnocline oscillating about the depths of the deployment site with a high-frequency interval, which might have come from internal waves breaking across the shelf. The absence of these high-frequency oscillations in our data does not imply that these waves do not exist but simply that we did not observe them during our short deployment, which

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was conducted during fair-weather conditions rather than a storm period when internal waves would most likely occur. Salinity and temperature measurements illustrate the general background conditions for this site. If our measurements are well and truly above or below the pycnocline, it could be that the signature of the internal waves (if present) are not reflected in the temperature and salinity signals but rather in the mean current flow, examined in the next section. Current Structure and Velocity Time-Series The key critical observations of mean current structure and time-series behavior are illustrated in Figures 7 and 8. The 300 kHz ADCP used in this study was set to the smallest vertical bin size (2 m) and a rapid profile repeat rate (30 second interval) in order to maximize the vertical and temporal resolution of the measurements with hopes of encountering high-frequency internal waves. Each 30-second burst represents an ensemble average of 16 individual acoustic pings over a vertical range from 4.2 m

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Figure 3. Side-by side comparison of side-scan sonar record showing raw georeference image (A) and geographic information system digitized bedforms (B). Star illustrates location of the acoustic Doppler current profiler/conductivity-temperature sensor /Rotary sonar bottom mooring.

Bedforms, coastal-trapped waves, and scour process observations above the bed to over 80 m above the bed, with horizontal and vertical measurements occurring every 2 m. A resulting color contour plot of the magnitude of the velocity (i.e., speed) of the horizontal current component, i.e., the Pythagorean addition of the east-west and north-south components of the flow, is illustrated in Figure 7. While a great deal of vertical and temporal structure exists in this figure, we will limit our discussion of the current velocity structure near to the bed (<10–20 m) especially the lowest bin at a height of 4.2 m above the bed, because the nearbed flows have the greatest influence on sediment transport and the formation of bedforms run-on. The lowest bin height is a function of the fundamental frequency of the ADCP (300 kHz), and the blanking distance beyond which valid velocity samples are returned. Of particular note are the two pulses of higher velocity near-bed flow that begin at ~5 h and 20 h respectively after the start of the deployment (Fig. 7). Each pulse episode lasts between 6 and 8 h with velocities ranging from 15 to 25 cm/s. The vertical structure during each pulse episode exhibits a general upsweep pattern in space and time with the highest velocities being recorded at the lower bins and decreas-

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ing away from the bottom up into the water column rising to as much as 30 m above the bed. Above this lower pulse layer is a zone 30–50 m above the bed where the speed drops to values of around 5 cm/s. This pattern suggests a velocity source from a lateral area moving across the mooring site and not a source from surface derived current (Fig. 7). Further insights into the local flow dynamics can be obtained by examining the time series of the lowermost bin. In Figure 8, the individual vector component time-series are portrayed for the bin 4.2 m above the bed. Here, several flow trends are evident. First, there is a deployment averaged net vector in the U component of 10 cm/s to the west and in the V component of 0.34 cm/s to the north. Overall, the current follows a progressive anticyclonic path generally following the shelf break toward the west with pulses toward the north-northwest. The two pulses toward the north-northwest are illustrated by the sharp bumps in the V component of the horizontal flow, with peaks in excess of 10 cm/s to the north (Fig. 8). The most important flows for sediment transport purposes are those with a competence (intensity) sufficient to initiate or

B

N bottom mount ADCP acoustic interference

superimposed Current vector

CT

rotary sonar

ADCP

Figure 4. (A) Schematic illustrating the configuration of the seabed mooring composed of acoustic Doppler current profiler (ADCP), conductivity-temperature sensor (CT) and rotary fanbeam sonar. (B) Rotary fanbeam sonar single pass scan image collected from the seabed mooring in the vicinity of the Chersonesos A site. High intensity backscatter returns are light pixels, while dark pixels are low backscatter.

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maintain the transport of sediment, in this case flows above the mean velocity of 12.45 cm/s. Based on the classic Hjulström curve (Hjulström, 1935) the mean flow (12.45 cm/s) encountered at the field site (Fig. 8) implies that the very fine to medium sands will remain in transport while the coarse sands and gravels exposed in the lower shell-hash/gravel facies will remain as un-entrained deposits. During peak observed flows (20 cm/s), the erosion

0hr

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threshold for the very fine and fine sands is met suggesting that fine/very-fine sands will be exhumed from the seabed exposing previously buried coarse sand/gravel shell hash units. This process of feedback between flow and scour sorting was previously documented in the work of Trembanis et al., (2004) and Green et al., (2004) for other shelf settings and suggests a similar set of processes are at work on the outer shelf of the Northern Black Sea.

+2.3hr

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Figure 5. Time-lapse images of rotary sonar scans over an 8 h period together with overlay of the mean current vector at the time of the scan (arrow). The time progression shows subtle ripples toward the outer ranges of the sonar scan but otherwise stasis in the bed configuration over this short time period. The current vectors illustrate a progressive anticyclonic mean current pattern with currents strengthening toward the S/SW.

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Salinity (PSS)

Temperature (C)

mean salinity: 20.6 psu

mean temperature: 8.3 °C

Hours from 8/15 2007 22:59 UTC

height above the bed (m)

velocity magnitude (cm/s)

Figure 6. Time-series plot of temperature (blue) and salinity (green) as recorded every 23 seconds by the conductivity-temperature sensor on the bottom mooring.

Time in hours since 8/15/07 19:59 UTC

Figure 7. Vertical cross-section color-contour plot of horizontal current speed (magnitude of velocity) measured between 4.2 and 84 m above the seabed in a local depth of 135 m. Bin 1 is located 4.2 m above the bed.

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TABLE 1. INSTRUMENT SAMPLING CONFIGURATION SUMMARY Depth No. of data Instrument Phenomenon (m) points Side-scan sonar Datasonics SIS-1000, Digitized bedforms over 113–237 2376 100 kHz, 200 m range, May 2006 entire shelf TRDI ADCP, 300 kHz, 2 m bins, Mean currents; transport 4.2–80 m above bed 11,021 August 2007 processes Imagenex 881-tilthead Rotary Sonar, Ripple geometry/evolution 135 (1 m above bed) 292 2.25 MHz, 9 m range, August 2007 SBE-37SM, August 2007 Salinity; temperature 135 (0.75 m above 6124 bed)

Energetic Near-Inertial Coastal-Trapped Waves The upper layer waters of the Black Sea are characterized by a predominantly cyclonic, rim current flowing along and occasionally across the continental shelf in a strongly variable spatiotemporal pattern. In addition, a series of quasi-permanent anticyclonic eddies exist inshore of the rim current (Oguz et al., 2005). The rim current structure is accompanied by coastaltrapped waves (CTWs) with an embedded train of eddies and meanders propagating cyclonically around the basin (Oguz and Besiktepe, 1999). According to the ADCP measurements

Sampling interval 1 survey

Sampling rate 1 survey

30 s 10 min

1 burst with 16 pings 3 scans

23 s

1 burst

(Oguz and Besiktepe, 1999), the rim current jet has a speed of 50–100 cm/s within the upper layer, and ~10–20 cm/s within the 150–300 m depth range. These shelf current features provide a mechanism for bidirectional transport between the nearshore and offshore regions. The most notable features of the Black Sea circulation system in relation to this study include (i) the meandering rim current system cyclonically encircling the basin and (ii) the Crimea and Sevastopol anticyclonic eddies on the coastal side of the rim current zone. It is important to note that the observations in this study were taken near the junction of the rim current and the Sevastopol anticyclonic eddy, termed a

Velocity (m/s)

V

U

mean U: -10 cm/s mean V: 0.34 cm/s Figure 8. Time-series plots of horizontal components of mean current measured every 30 seconds at bin 1 (4.2 m above the bed). Red line shows north-south flow (north positive) and the blue line represents east-west flow (east positive).

Bedforms, coastal-trapped waves, and scour process observations “quasi-permanent or recurrent feature” in the upper layer circulation by Oguz et al. (2005). Because of strong stratification in the Black Sea, CTWs reach the inertial frequency. Furthermore, their velocity fluctuations are surface-intensified off the shelf break (over the upper slope), but reverse to bottom-intensified on the shelf (Ivanov and Yankovsky, 1993). The velocity amplitude of CTWs on the South Crimean shelf is ~10–15 cm/s in the summer and is likely to be higher during the winter season notorious for its severe storms. CTWs propagate with the coast on the right (as a Kelvin wave), westward in the Northern Black Sea. As CTWs encounter coastline and topographic variations past Cape Sarych (the southernmost tip of the Crimea Peninsula) they scatter into other (higher) wave modes in order to adjust to these “disturbances” of the waveguide. Scattering typically introduces smaller spatial scales and higher amplitudes in the wave field. Indeed, the outer shelf and slope just to the west of Crimea, in the vicinity of our deployment site (star in Figs. 1–3), is characterized by strong mesoscale variability with frequent predominantly anticyclonic eddies. This particular feature of the regional shelf dynamics prompted a series of papers aimed at examining the phenomenon (i.e., Yankovsky and Chapman, 1995, 1996, 1997). In these papers, the authors found that very energetic vortex-like current patterns could be developed in the near bottom layer with near-inertial or subinertial frequencies. In addition to the CTWs, the rim current itself meanders and makes episodic onshore excursions. An example of this behavior can be found in the vertical transect of the rim current by Oguz and Besiktepe (1999; their fig. 3). The implications of these processes (CTWs, meanders in the rim current, and potential storminduced internal waves) is that there are a number of physical processes that can promote cross-shelf transport (including anoxic waters) across the shelf and contribute to the scour and preservation potential in the region. Bedform Distribution and Scour Patterns During the 2006 field campaign, over 300 km of linear survey line was run with the 100 kHz side-scan sonar towfish. Each survey line was processed for water-column removal, layback positioning, and beam angle correction using the SonarWiz™ (Chesapeake Technology Inc.) processing system to generate a final georeferenced sonar mosaic as illustrated in Figure 2 (inset). Next, each line was examined and all discernable bedforms were digitized (Fig. 3) using the heads up GIS feature of SonarWiz. In total, 2376 individual bedforms were digitized in this manner providing a GIS data set for subsequent analysis. Both the sidescan sonar mosaic and the digitized bedforms were exported to the Google Earth™ KML file format to allow viewing and wider distribution of the results (Figs. 2 and 3). Once digitized, wavelength spacing between each successive bedform was calculated from the centroid of each bedform to the next, with a resulting mean and modal wavelength calculation of 72.8 and 15.7 m respectively (Fig. 9A). Furthermore, for each bedform, a corre-

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sponding strike orientation (along-crest) and dip (crest orthogonal) orientation was determined (Fig. 9B). Recall that azimuth and dip are orthogonal, therefore, the principal dip orientation as determined from least squares rotation analysis was found to be 52°/232° (Fig. 9B). The 180° ambiguities in bedform dip direction (52°/232°) are a result of the two-dimensional nature of side-scan sonar data and the lack of bathymetric data to determine the stoss (updrift) face from the lee (downdrift) face of the bedform. In Figure 9C, a comparison between the bedform dip orientation (blue dots) and the nearbed mean current (velocity amplitude of 12–20 cm/s) displayed a close association with peak flows (red arrows) in a crest-orthogonal direction. It seems most likely given the observed flow regime that the direction of lee face is 232°. A progressive vector diagram of the nearbed current (Fig. 10) also shows the connection between bedform orientation and shelf transport flows. Assuming an idealized homogeneous lateral flow regime, the progressive vector diagram illustrates the path that a particle of water near the bed would take over the period of the bottom mount deployment. This shows the general westerly flow with several turns to the north-northwest associated with the strong pulses of nearbed flow, further emphasizing that adiabathic flows do commonly occur in this region. SUMMARY AND CONCLUSIONS The Northern Black Sea shelf is a complex, heterogeneous, morphodynamic system that provides a critical link between the nearshore and the deep sea (Fig. 1). Marked spatial and temporal variations in bedform geometry as well as sharp gradients in forcing conditions (rim current, CTWs) typify the shelf setting. Increasing our present understanding of the coupled links between variable bed geometry and hydrodynamics is a crucial step in order to decipher the morphologic history of the Black Sea shelf since the end of the Last Glacial Maximum. In this study, we found no appreciable variation in nearbed salinity or temperature during the deployment of our instrumented bottom mount, suggesting absence of baroclinic driven flows during this period (Fig. 5). Furthermore, the ADCP measurements of nearbed flows also gave no indication of high-frequency internal wave activity during the deployment, suggesting that these events may be highly seasonal and event driven. Nevertheless, significant and distinct nearbed flow features were observed during the field deployment as evidenced by the observed quasi-steady, near-inertial flows that were decoupled from surface and showed strongest velocities near the bed (Figs. 6 and 7). The currents, possibly CTWs oscillated with ~15 h periodicity and amplitude of ~15 cm/s. Previous physical oceanographic work in the region suggests that there are several reasons to expect strong variations of near bottom currents in the near-inertial–subinertial frequency range on the outer shelf southwest of Sevastopol including CTWs and meanders in the rim current. This study presents a new observational investigation of bedform distribution and the relation of bedform geometry to

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nearbed hydrodynamics (Fig. 8C). Previously suggested theories suggest that these bedforms are static ancient features, a theory that is perhaps correct only for isolated instances. The observed flow dynamics and bedform orientations imply an increased preservation potential of artifacts on the shelf. Furthermore, the observed outer shelf flows (Fig. 10) may also reduce the depth of viable preservation in artifacts that would otherwise exist above the anoxic layer. The occurrence of presently active bedforms on the outer shelf as documented here, further emphasizes the role that CTWs, internal waves, and rim currents have on the transport of water and sediment across the shelf.

ACKNOWLEDGMENTS The authors wish to acknowledge the financial and technical support of the following groups: the officers and crew of the NRV Alliance, the office of the President of Ukraine, U.S. National Oceanic and Atmospheric Administration (NOAA) Office of Ocean Exploration and Research, the Office of Naval Research, the Center for Coastal and Ocean Mapping at the University of New Hampshire, and the National Geographic Society. The authors particularly thank Drs. Tony Rodriguez and Doug Levin for reviews that substantially improved the paper.

Number of Occurrences

A

C

bedform wave-length (m)

B

Legend arrows : mean currents dots : bedform wave ray

Bedform orientation (dip direction) Figure 9. (A) Histogram of bedform wavelength (e.g., spacing between crests) for each of the bedform pairs in the digitized record. N = 2376 and mean wavelength 72.8 m (B) Compass plot of bedform of dip azimuth (orthogonal to crest), calculated from the geographic information system digitized side-scan sonar. N = 2376 and principle dip azimuth of 52/232 degrees true north. Crest orthogonal azimuth has a 180° ambiguity depending on which side of the bedform is used for the origin reference, therefore the conjugate pair for each crest orthogonal is shown (blue arrows) and the principle bedform orientation is shown by the long red arrow. (C) Mean currents (red arrows) overlain on bedform wave ray (blue dots). Principle bedform dip orientation, determined by least squares rotation analysis, is shown by the thick red arrow. Note: Currents are in the direction of propagation and bedforms are shown as conjugate direction across crest.

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60 km

Sevastopol progressive current vector

Digitized bedforms 10 km Figure 10. Progressive vector diagram plot of bottom acoustic Doppler current profiler bin mean current illustrating net current transport near the bed.

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The Geological Society of America Special Paper 473 2011

Archaeological oceanography and environmental characterization of shipwrecks in the Black Sea Michael L. Brennan Robert D. Ballard Katherine L. Croff Bell Graduate School of Oceanography, University of Rhode Island, South Ferry Road, Narragansett, Rhode Island 02882, USA Dennis Piechota Fiske Center for Archaeological Research, University of Massachusetts, Boston, Massachusetts 02125, USA

ABSTRACT The August 2007 expedition to the Black Sea continued a multiyear project designed to locate and study ancient shipwrecks in deep water. The expedition revisited and investigated two shipwrecks, Sinop D (at 325 m depth off Sinop, Turkey) and Chersonesos A (at 135 m depth off Sevastopol, Ukraine). These wreck sites are good case studies for our research because they are located in different parts of the Black Sea, in anoxic and suboxic waters, respectively. Preliminary data reported here are from seawater samples taken from around the wrecks and a year-long collection of temperature, salinity, and pressure data. Trace-element data from the seawater samples are consistent with reported processes and values for the Black Sea. The oceanographic sensor data confirm the stagnant nature of the anoxic water layer that has allowed for the high level of preservation of the Sinop D wreck site. We also discuss the design and placement of two sets of experiments left in situ to characterize the decay rates of common materials found on ancient shipwrecks, including wood and metal. By providing ways to understand the chemical and physical processes that characterize different parts of the Black Sea water column, these wrecks are important sites for (1) determining the preservation potentials of cultural materials in deep water, and (2) informing the design of methodologies necessary to conserve them.

INTRODUCTION This paper discusses some preliminary results and the future direction of a multiyear program in the Black Sea initiated with an August 2007 expedition aboard the R/V Alliance. This expedition was a continuation of work begun by the Institute for Exploration (IFE) in 1998 and continued in conjunction with the University

of Rhode Island’s (URI) Center for Ocean Exploration (COE) in 2006. Our work in the Black Sea serves as an important case study that illustrates the way we approach the identification and documentation of archaeological sites in the deep sea. This complex and largely unexplored environment in which archaeological materials have been deposited on the seafloor as shipwrecks demands an approach different from that of traditional nautical

Brennan, M.L., Ballard, R.D., Croff Bell, K.L., and Piechota, D., 2011, Archaeological oceanography and environmental characterization of shipwrecks in the Black Sea, in Buynevich, I.V., Yanko-Hombach, V., Gilbert, A., and Martin, R.E., eds., Geology and Geoarchaeology of the Black Sea Region: Beyond the Flood Hypothesis: Geological Society of America Special Paper 473, p. 179–188, doi: 10.1130/2011.2473(11). For permission to copy, contact [email protected]. © 2011 The Geological Society of America. All rights reserved.

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archaeology. The methods by which we locate, investigate, and conserve these wreck sites require detailed knowledge of the environmental conditions in which the wrecks lie, as well as an understanding of how these alien materials interact with their environment (e.g., Palma, 2005). Archaeology has become increasingly reliant on the application of scientific techniques for the analysis of cultural materials. As the capabilities of archaeological science continue to expand, so too does our responsibility to survey, document, and conserve these sites appropriately. Ballard (2008) pointed out that oceanography is not a stand-alone discipline but incorporates other scientific disciplines, and it should not be difficult to incorporate archaeology into its field of research. Archaeological oceanography utilizes oceanographic, geological, chemical, biological, and physical methods to investigate and interpret cultural material deposited on the seafloor. Beyond the oceanographic methods used in archaeological surveys described by Coleman and Ballard (2008), archaeological oceanography also aims to use shipwrecks as platforms for environmental studies of the surrounding ocean. The occupational cycle of an archaeological site ends with its abandonment, however, we are also concerned with understanding the postabandonment environmental history of these sites as much as we are with the history of the site’s origin and use. In the case of a shipwreck, this postabandonment history begins when the ship sinks and is deposited on the seafloor. Through chemical and physical monitoring, and geochemical analytical methods, we can use the material from and around a wreck site to learn about the submerged and buried environment, i.e., its repository. In essence, we use oceanography to investigate archaeological sites and their constituent materials to increase our understanding of the oceanographic setting. The Black Sea is an ideal location for the development and implementation of an archaeological oceanographic methodology. The highly stratified water column and anoxic bottom waters allow for a high level of preservation of organic material from cultural sites, redefining the ways in which we need to approach archaeological materials from shipwrecks. The oxic-suboxic-anoxic layered structure of the Black Sea creates a special opportunity to study archaeological materials in a variety of oceanographic and sedimentary environments and conduct comparative analysis of the resulting effects on such materials. The Black Sea thus constitutes a natural laboratory in which the different water layers can be studied in a single water column, making it an excellent case study for discussing archaeological oceanography—its theory, methods, and potential for combining the efforts of science and archaeology in order to understand the dynamics of an archaeological and environmental system. BACKGROUND Black Sea Environments The Black Sea is the largest anoxic basin on Earth, formed because of the single narrow and shallow outflow channel of the

Bosporus, which has caused the sea to become highly stratified. The flow and circulation patterns in the Black Sea are generally controlled by its limited exchange of water with the Mediterranean Sea through the Bosporus Strait and into the Aegean Sea (Ross, 1977). Below the permanent pycnocline at ~150 m depth, the waters of the Black Sea become anoxic. This stratification of the water column is a consequence of the density differences between the inflowing fresh river water and the bottom current from the highly saline Mediterranean. The Black Sea’s two-layer current system is the result of estuarine circulation caused by a positive water balance due to riverine influx (Lyons, 1992). Shipwrecks in this body of water provide a good opportunity to find organic material in various states of preservation not often seen in equivalent land or shallow-water sites. They also serve as fixed points for comparative studies of the interaction of the Black Sea’s sediments and water layers with cultural materials. As a function of their contrasting densities, the waters of the Black Sea are vertically stratified into three distinct layers: an upper oxic layer, a bottom anoxic layer, and a transitional suboxic layer that was first identified by Murray et al. (1989). The suboxic zone is defined as the region of the water column where oxygen has decreased to <10 µM and where hydrogen sulfide (H2S) has not yet appeared in a concentration greater than 10 µM. In response to climate changes, the thickness of the suboxic layer is believed to have varied over time (Murray et al., 2007). More importantly, the suboxic zone is a barrier to the coexistence of oxygen and sulfide in the Black Sea water column (Glazer et al., 2006). The depth of the anoxic zone is less variable than that of the oxycline. Therefore, the depth and thickness of the suboxic zone is governed by the ventilation of the thermocline by oxygenated surface waters and the oxidation of organic matter (Murray et al., 2007). Given that dissolved oxygen and sulfide do not coexist between the oxic and anoxic layers, the mechanism for sulfide oxidation in the anoxic zone cannot be a downward flux of oxygen from the surface waters. Several studies have indicated that dissolved and particulate manganese (Mn) may be the sink for the upward flux of H2S from the anoxic layer, since there is no downward flux of oxygen (Luther et al., 1991; Millero, 1991; Murray et al., 2007; Trouwborst et al., 2006). Byzantine Shipwrecks Shipwrecks can be used as fixed platforms from which longterm monitoring of oceanographic conditions can be conducted. IFE and COE have been studying the coastal geology and submerged cultural landscape of the Black Sea since 1998 and have documented numerous ancient shipwrecks off the coasts of Bulgaria, Turkey, and Ukraine, dating back as early as the Hellenistic period (Coleman and Ballard, 2007). Recent expeditions off the coasts of Ukraine and Turkey have begun investigating two Byzantine shipwrecks in more detail through environmental monitoring, site mapping, and localized excavations with the remotely operated vehicle (ROV) Hercules. These wrecks,

Oceanography and environmental characterization of shipwrecks named Chersonesos A and Sinop D for the ports near which they were discovered, were found in significant positions on the Black Sea bottom, and they present an ideal opportunity for a comparison between water layers across the oxic-anoxic boundary. Chersonesos A is located at 135 m depth at the edge of the shelf break, while Sinop D lies at 325 m depth along the continental slope (Fig. 1). These shipwrecks are positioned on the margin of the oxicanoxic interface and are fixed points in the Black Sea from which a comparison between the anoxic and suboxic waters can be conducted on similar cultural materials. This paper discusses the methods with which these two wrecks have been and will continue to be studied in the interest of understanding the different environments of the Black Sea’s stratified water column. These ships from similar time periods can be considered to have contained and been constructed of similar materials. Therefore, by first understanding the two environments in which each wreck lies, we can interpret the current state of the wrecks and better investigate and conserve these sites.

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Sinop D Four shipwrecks were identified in 2000 off the coast of Sinop, Turkey (Fig. 1) and were further investigated in 2003. Wrecks Sinop A, B, and C were found between 85 and 101 m depth and are ovoid mounds of carrot-shaped amphorae, a ceramic type commonly associated with Sinopean production (Ward and Ballard, 2004). The fourth wreck, Sinop D, is located at a depth of 325 m, fully within the anoxic layer, and it is very well preserved, with its mast and vertical posts still standing (Fig. 2). The sediment at this depth is deep and very soft; the ship appears to have sunk into the sediment, leaving only the mast and frame ends of the ship exposed. Localized excavations were carried out in 2003 and 2007 to remove the sediment that partially buries the wreck and all of its cargo. Three amphorae were recovered in 2003, and an additional one was moved to an offsite storage rack in 2007 from the midships area. Radiocarbon dating and ceramic type analysis have placed this shipwreck in the fifth century A.D. (Ward and Ballard, 2004; Ward and Horlings, 2008). The limited scope of the excavations to date has revealed

Figure 1. Map of the Black Sea showing the locations of the shipwrecks Chersonesos A and Sinop D. Bold line marks the 150 m depth contour and approximate anoxic boundary. Bathymetry data are from the GEBCO_08 Grid, version 20091120 (http://www.gebco.net).

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only a small portion of the wreck’s cargo, however, and little can be said with certainty about the hull construction. Chersonesos A An additional Byzantine shipwreck was discovered in 2006 ~23 km off the southwest coast of Crimea, Ukraine, at a depth of 135 m (Fig. 1); typological ceramic dating places it between the ninth and eleventh centuries A.D. (Andrei Opait, 2006, personal commun.). The wreck is poorly preserved by comparison to Sinop D, but some timber and hull remains were evident before excavation (Fig. 3). Oxygen readings from the sensor on the ROV Hercules showed that dissolved oxygen was present in concentrations less than 5 µM, placing the wreck in the suboxic zone. The seafloor on the shelf at this depth is a harder surface than that of the Sinop D wreck site, and the minimal settling of the ship into sediment has provided less support for holding the wreck together. The cargo of Chersonesos A consists primarily of flatbottomed, one-handled jars. These artifacts are grouped into two mounds, with sediments and modern debris built up around the wreck site (Buxton et al., 2008). Although remnants of the ship’s timbers are still evident, its frame has clearly deteriorated to a much more decomposed state than that of Sinop D. METHODS AND RESULTS The focus of the 2007 expedition was the implementation of a number of multiyear environmental monitoring experiments and data collection at the locations of Chersonesos A and Sinop D that will provide a broad understanding of the long-term processes at these sites. The intention is to use these two fixed locations as a means of investigating the environments across the suboxic-anoxic interface. The benefit of utilizing these two archaeological sites as platforms for oceanographic research is our ability to leave sensors in place at specific points and to return

Figure 2. Photograph of the Sinop D wreck site at 325 m depth off Sinop, Turkey. The view looks obliquely downward toward the northeast from the stern area of the wreck. Note the well-preserved frame timbers and mast. ©IFE/COE; used with permission.

to the same location over multiple years to collect redundant samples for comparison. Some preliminary data were collected in 2007 that allow for initial results to be discussed. At both the Sinop D and Chersonesos A wreck sites, sediment cores, CTD (conductivity-temperature-depth), dissolved oxygen measurements, and water samples were collected by Hercules while on site. Each wreck was also photographed and mapped periodically during the ROV’s work, documenting the progress of the excavations as additional timbers and cargo were exposed. Two ceramic jars (given artifact codes AAB and AAJ) from the Chersonesos A cargo were recovered at the request of the state of Ukraine and were conserved there. The exterior flat bottoms of both jars contain a black sandy material that appears to have been pressed into the clay prior to firing. Our petrographic analysis of a comparable artifact, a jar from the terrestrial site of Crimean Chersonesos provided by Andrei Opait from the Institute for Classical Archaeology, showed the presence of a grog and feldspar-rich temper as well as large amounts of iron oxide, but it did not contain the black sand grains on the bottom. Framboidal pyrite (FeS2) was found in sediment collected from inside jar AAB during conservation (Fig. 4). Additionally, a set of experiments that was left near the two wreck sites to test the redox and decay rates of various organic and metal samples will be recovered during later expeditions. Environmental Monitoring Two sensor packages were placed off the coast of Sinop, Turkey, at the end of the expedition in August 2007. The first was equipped with a Seabird CTD and current meter; it was deployed near the Sinop D wreck site at a depth of 330 m. The second was equipped with an Aanderaa CTD and optode, and RDI ADCP current profiler; it was placed 8.5 km southeast of Sinop D at a depth of 106 m within the suboxic zone. The deep sensor package

Figure 3. Photograph of the Chersonesos A wreck site at 135 m depth off Sevastopol, Ukraine. The view looks obliquely downward toward the south. Note the prominent cargo of partially buried flat-bottomed ceramic jars and absence of preserved frame timbers. ©IFE/COE; used with permission.

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0.0136 °C, 21.74 ± 0.0441 psu, 456.646 ± 0.1346 psi, and 1016.8 ± 0.033 kg m–3, respectively. The correlation coefficient between temperature and salinity is 0.825. Each of these data sets shows a very small variation from the mean (Fig. 5). Mass Spectrometry of Bottom Water Samples

Figure 4. Photomicrograph of framboidal pyrite recovered from sediment inside ceramic jar AAB from Chersonesos A. ©IFE/ COE; used with permission. Photograph by Dennis Piechota.

was recovered in September 2008. The sensor package deployed in the suboxic zone was not located; efforts will be made in subsequent years to re-locate and recover this sensor package. Data were logged from 25 August 2007 to 28 September 2008. Standard deviations were calculated for temperature, salinity, pressure, and potential density with values of 8.809 ±

One water sample was collected each from the oxic, suboxic, and anoxic water layers using a Niskin bottle attached to Hercules. The oxic sample was taken from bottom water of the Sea of Crete to serve as a control. The two other samples were collected during dives on Chersonesos A and Sinop D. All three bottom water samples were taken within 5 m of the seafloor. The water samples were filtered with 0.45 µm filters, acidified to 2% with nitric acid (HNO3), and diluted by a factor of 0.1. They were then analyzed using a Thermo X-Series II quadruple ICPMS (inductively coupled plasma–mass spectrometer) at URI’s Graduate School of Oceanography following analytical methods outlined in Kelley et al. (2003). We also analyzed a multi-element solution for calibration and two blanks.

Figure 5. Graph of conductivity-temperature-depth sensor data from 25 August 2007 to 28 September 2008 showing temperature (°C), salinity (ppt), pressure (dbar), and potential density (kg m–3) versus time.

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The water sample collected by Hercules from the Sea of Crete was analyzed along with the Black Sea samples from the two wreck sites as a control for oxygenated bottom water. The analytical data from this sample show a general correlation with the data from the Chersonesos A suboxic water sample (Fig. 6). The major differences among the three samples are in the elements vanadium (V) and manganese (Mn) of the Sinop D anoxic sample. Rare earth elements (REE) were mostly present in concentrations just at or below detection levels. These data will not be included in the analysis. DISCUSSION Environmental Monitoring The temperature, salinity, and pressure data from the CTD sensor package moored at the Sinop D wreck site illustrate the stagnant nature of the anoxic water column. The standard deviations reported here show little variation around the Sinop D wreck site. Figure 5 illustrates the 13 mo data set of temperature, salinity, pressure, and potential density versus time. Temperature, salinity, and density show similar trends, with a good correlation between temperature and salinity. The pressure data do not correlate well with temperature or salinity. The small variations in each of these data sets could be reflective of seasonal atmospheric changes. A large storm crossed the Black Sea on 10 November 2007 with 45 knot winds and 5–7 m wave heights (Weathernews Inc., 2007). The disturbance caused by this low-pressure system is visible in the data set as a downward spike in pressure. Temperature, salinity, and potential density exhibit a downward spike a few days later, which may reflect an effect of the storm on the water column (Fig. 5). While this disturbance is apparent in the data, it caused only a very small variance at a depth of 330 m. Generally, these data confirm our observations during the expedi-

tion that the anoxic water column exhibits little variability over the time scale of a year. Bottom Water Chemistry The trace-element data from the bottom waters at the two wreck sites are restricted to a single sample from each site. Analyses of water from these two locations illustrate the presence of the metals and trace elements that are in solution at each wreck site, one anoxic and the other suboxic. The dissolved trace-element comparison between the suboxic and anoxic waters allows us to understand the likely mineralogy of authigenic components in the sediment cores recovered from the wreck sites. An interesting question for future work is whether episodic internal waves pass over the Chersonesos A wreck site, as Trembanis et al. (this volume) suggest, and whether they influence suboxic water chemistry to the extent of affecting precipitates in the sediment. Iron plays an important role in anoxic water as a sink for H2S in the form of sedimentary pyrite (Berner, 1970). The role of Mn is less well understood, but recent research has shown that it may serve an equally important role as MnO2, which oxidizes an upward flux of H2S at the suboxic-anoxic boundary (Murray et al., 2007). Trouwborst et al. (2006) showed that the Mn(III) chemical species can act as an oxidant or reductant and is the species that maintains the existence of suboxic zones, allowing for exchange between oxic and anoxic zones without the coexistence of oxygen and hydrogen sulfide. In anoxic environments, Fe and Mn are present in their reduced forms (Fe[II] and Mn[II]) as metal ions (Fe2+ and Mn2+), as opposed to oxic waters, in which they exist as metal oxyhydroxides in their oxidized forms, Fe(III) and Mn(IV) (Balistrieri et al., 1994). The change in chemical species occurs within the suboxic zone. Further reactions occur below the suboxic region, in the anoxic zone. Here, Fe and Mn behave slightly differently. Fe,

Figure 6. Graph of trace-element data from samples from the Sea of Crete, and the Sinop D and Chersonesos A shipwreck sites.

Oceanography and environmental characterization of shipwrecks for example, oxidizes faster (Balistrieri et al., 1994). More importantly, Fe plays a much greater role in the sulfur cycle than does Mn, and iron sulfides form directly from interactions between dissolved ferrous iron and sulfide. This can separate Fe(II) and Mn(II) at redox boundaries, as shown by Lewis and Landing (1991). Iron sulfides formed in the anoxic region adsorb other trace metals, including Co, Cu, Mn, and Mo, and are an important sink for them in the form of authigenic pyrite (Balistrieri et al., 1994). Our analysis of water samples from Chersonesos A and Sinop D in the suboxic and anoxic regions, respectively, showed surprisingly similar values between wreck sites for most trace metals, including Fe, Co, Ni, Cu, and Mo (Fig. 6). Such results may suggest that these elements have been effectively removed from the bottom water by pyrite formation. The elements that showed the greatest differences between the suboxic and anoxic water samples were V and Mn. Both had much higher concentrations in the anoxic sample from Sinop D. Many trace metals in anoxic environments (e.g., Mo, Ni, Cu, and Zn) commonly adsorb onto pyrite mineral surfaces, which act as a sink for them, transporting them to the sediment (Tribovillard et al., 2006; Vorlicek et al., 2004). Vanadium is coupled with the redox cycle of both Mn and Fe and adsorbs onto oxyhydroxide structures. However, it is not taken into pyrite or other Fe-sulfides like many other trace metals (Tribovillard et al., 2006). Similarly, dissolved Mn is not significantly substituted into Fe-sulfide phases. Because of its cycle across redox boundaries by oxidizing H2S, dissolved Mn concentrations are high just below the suboxic-anoxic interface. Reduced Mn-oxyhydroxides settle through the anoxic water and release adsorbed trace metals, which can then be adsorbed onto pyrite near the sediment-water interface (Huerta-Diaz and Morse, 1992; Tribovillard et al., 2006). Therefore, when comparing the samples of bottom water from Chersonesos A (suboxic at 135 m) and Sinop D (anoxic at 325 m), it appears that the role of the Mn redox cycle across the suboxic-anoxic interface and the formation of pyrite in the anoxic bottom waters are both very important in the cycling of dissolved trace metals. Molybdenum and other metals that substitute for Fe in pyrite exhibit a pattern similar to the Fe concentrations in the water samples, which was similar between the suboxic and anoxic samples. Manganese and V, two elements that do not significantly substitute for Fe in the formation of authigenic pyrite, showed greater concentrations in the anoxic water, as Fe-sulfide formation would not have removed them to the sediment. A preliminary analysis of core C1003–06 from near the Sinop D wreck site with a Niton XL3t 900 X-ray fluorescence (XRF) analyzer showed that Fe concentrations are much higher in the sediments than Mn, while Mn is greater in the anoxic water sample due to the removal of Fe through authigenic pyrite formation. These results will support mineralogical analyses of the sediment cores from the wreck sites and help in a comparison between those from the Chersonesos A and Sinop D locations.

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LONG-TERM STUDIES AND FUTURE WORK Sediment Cores In total, 12 push cores were recovered at the two wreck sites by Hercules, four from Chersonesos A and eight from Sinop D. Two of those obtained from Chersonesos A were taken 10 m away from the wreck and the others from directly adjacent to the exposed frames of the hull. One core from Sinop D was recovered ~5 m away from the wreck and the others from points around the perimeter of the wreck, just outside the exposed timbers. The last two cores were taken from an excavated area inside the exposed frames of the ship. Two of the cores were split, and parts were sampled for biology because possible bioturbation was observed inside them. A variety of analyses are currently being conducted on the cores, including X-rays, micromorphology, bulk properties analysis, XRF elemental analysis, and grain-size analysis. These cores will allow us to compare the sediments around the two wrecks, and they will also show how the wreck sites interact with the seafloor by revealing morphological and mineralogical differences between cores taken right at the wrecks and at distances from them. Mineralogical and trace-element data from the cores will be correlated with data from the water samples and sediment recovered from inside the jars in order to understand the chemical processes at each site. In Situ Experiments Two types of environmental experiments were designed and constructed by the conservator Dennis Piechota and mechanical engineer Todd Gregory. They were left in place at the Sinop D and Chersonesos A wreck sites. The objectives of deploying these ancient material analogues as decay characterization experiments are twofold: (1) to assess the character and strength of the natural deteriorative forces experienced by each wreck and their artifacts over centuries of submergence, and (2) to assess the suitability of the open-water environment at each site for the underwater storage and display of artifacts in future years. At each site, we deployed two types of ongoing decay rate experiments called “kebabs” (Fig. 7) and “twinkies” (Fig. 8). In these tests, modern materials comprising samples of wood, metal, bone, grain, and rawhide are exposed to the open water and sediment environments simultaneously. After incremental test periods, these samples will be retrieved for analysis to infer the rates and types of decay currently acting above and below the seabed. The “kebab” samples were made into small (2.5 cm diameter × 2 cm height) cylindrical coupons of wood and carbon steel and slid onto an internal titanium rod to form a stack (Fig. 7). The rod with its test samples was deployed as a spike driven vertically into the sediment by Hercules to record the physical, chemical, and biochemical effects the seawater and sediment have on these surrogates of cultural materials commonly found on ancient wrecks. Two types of kebabs were deployed: wood and metal. The wood kebab carries alternating samples of oak and pine. Besides being commonly used in shipbuilding, these samples

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represent the two primary forms of wood microstructure: ringporous (hardwood) and diffuse-porous (softwood). A second type of kebab was arrayed with cylinders of a modern mild steel (type 1020) as indicators of the rates and types of corrosion that will be experienced by metals in open seawater and at varying sediment depths. Eight of these spikes, four containing steel samples and four wood, were each deployed at Chersonesos A and Sinop D. Excavation schedule permitting, the first pair of spikes, one wood and one metal, will be retrieved in 2011 for deterioration analysis, and the remaining pairs will be retrieved in 2015, 2019, and 2023. The spike design will allow us to record how the decay and redox of wood and metal materials change with sediment depth between suboxic and anoxic environments. Most underwater archaeological sites are exposed to oxygenated waters at the seabed interface. This leads to the complete loss of organics and rapid corrosion rate of metals. Preservation at such sites relies on burial of cultural remains in sediments deep enough to create and maintain anoxic conditions. Increased preservation of cultural remains is normally visible during excavation down from the sediment-water interface and into anoxic sediments. At these two Black Sea sites, we expect good preservation of cultural materials to occur from the sediment-water interface down into the sediments as well as to a certain degree above the interface, especially at Sinop D. These kebab experiments will indicate the differences in decay rates and decay types between the suboxic site of Chersonesos A and the anoxic site of Sinop D. They will also show whether deeply buried components of ships, whether metal or wood, can be safely stored after excavation in the open water of the seabed adjacent to these Black Sea wreck sites. A second type of decay rate platform was designed and deployed at each site to test additional materials. These “twinkies” are open boxes that were designed simply to sit on the seafloor (Fig. 8). Each twinkie contains duplicate samples arranged to face the open water and sediment surface simultaneously (Fig. 9). The materials in these experiments include wood

Figure 7. Photograph of “kebabs” loaded with test samples of: (A) wood (alternating samples of oak and pine); and (B) mild steel. ©IFE/COE; used with permission. Photograph by Dennis Piechota.

and steel samples similar to those used in the kebabs; barley, a proxy for a common historic cargo; rawhide and bone samples to begin addressing whether human remains may be preserved; and additional historic metal analogs of copper and lead. Because of technical problems at Chersonesos A, we were able to deploy only two of the four twinkies for that site. All four twinkies were deployed at the Sinop D site. These tests will be recovered in the coming years at the same time as the kebabs to begin determining the decay rates and processes for these sample materials. Once retrieved, each sample will be immediately photodocumented for overall condition. They will then be subsampled for multiple analyses by specialists in bacteriology and elemental analysis. The wood and proteinaceous samples will be examined to determine the types and extent of microbial attack, the degree of wood hydration (waterlogging), and overall microphysical alteration. Metal samples will be examined to determine the total weight loss of the eutectic material (remaining uncorroded metal) and to identify the corrosion products and their crystalline forms using microscopy and instrumental analysis. The results of these experiments will allow us to characterize the differences in decay rates at the two wreck sites, both biological and chemical, and both above and below the sedimentwater interface (e.g., Björdal and Nilsson, 2008; Palma, 2005). These data, along with the data from sediment cores and water samples, will help us to define and compare the environments in which these two ancient shipwrecks lie. CONCLUSIONS Our work at these two wreck sites in the Black Sea is still in its early stages. The preliminary data discussed here establish a framework for further research on subsequent expeditions. While these Byzantine wrecks are important historical finds due to the rare discovery of ancient wrecks in the deep sea, their greater

Figure 8. Photograph of “twinkie.” The water-exposed samples are arranged on panels within the crate, while the sediment-exposed samples are mounted on a separate panel under the Neoprene skirt. ©IFE/COE; used with permission. Photograph by Dennis Piechota.

Oceanography and environmental characterization of shipwrecks significance comes from their locations within the Black Sea and the potential for a unique study. The comparison of the Sinop D and Chersonesos A wreck sites goes beyond the circumstances of their respective preservation in the anoxic and suboxic layers of the water column. Rather than limit our studies to the physical elements of the wrecks themselves, archaeological oceanographic methodology works to characterize the environmental setting of the wrecks by utilizing their fixed locations as platforms for longterm studies. The environmental monitoring, water sample analyses, and long-term experiments we report here are merely the first step in this process. We plan to investigate these wrecks further to understand more fully the dynamics between the environments and cultural materials, as well as the effects each has on the other. The Black Sea expedition in August 2007 was the first time this particular application of environmental monitoring and in situ preservation has ever been attempted at a deep-water site. In recognition of the new standards of practice advocated in the Annex to the 2001 UNESCO Convention on the Protection of the Underwater Cultural Heritage, which encourages this in situ

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research as a preferable strategy to full excavation, our work in the Black Sea aims to define the effects of the anoxic and suboxic environments on cultural materials. This research will hopefully pave the way for onsite storage of excavated artifacts, especially at the Sinop D wreck site. Consequently, the Black Sea constitutes a natural laboratory for the design and implementation of archaeological oceanography methodology. The use of these two shipwrecks as platforms to study the surrounding environment will facilitate their future investigation and conservation in a more complete, yet less destructive, manner. By characterizing the marine environments in which shipwreck sites have formed, we approach the wrecks as both historical media and modern marine sites, at which active processes continually interact with the cultural materials in complex dynamic systems. Differences in these systems between wreck sites composed of similar materials, as in the case of the Byzantine Sinop D and Chersonesos A wrecks, provide a unique opportunity for comparisons of these marine settings. In this way, archaeological oceanography goes beyond marine archaeology and oceanographic survey to combine the disciplines and allow us to take a new approach to deep-water archaeology. ACKNOWLEDGMENTS The authors thank Stuart Bishop, Bridget Buxton, Ilya Buynevich, Kathleen Cantner, Alexis Catsambis, Dwight Coleman, Dan Davis, Todd Gregory, Anastasiya Iyevlyeva, Katherine Kelley, Eric Martin, Jim Newman, Andrei Opait, Brennan Phillips, Webb Pinner, Chris Roman, Sergiy Voronov, Mark Wimbush, the Department of Underwater Heritage, Academy of Sciences of Ukraine, and the captain and crew of the R/V Alliance. This research project was funded through a grant from the National Oceanic and Atmospheric Administration Office of Ocean Exploration. REFERENCES CITED

Figure 9. Photograph of a panel of “coupons” of various test materials deployed on the underside of a twinkie. When the twinkie is deployed on the seafloor, the coupons are in contact with the sediment and separated from exposure to open water by a Neoprene membrane. ©IFE/ COE; used with permission. Photograph by Dennis Piechota.

Balistrieri, L.S., Murray, J.W., and Paul, B., 1994, The geochemical cycling of trace elements in a biogenic meromictic lake: Geochimica et Cosmochimica Acta, v. 58, p. 3993–4008, doi: 10.1016/0016-7037(94)90262-3. Ballard, R.D., 2008, Introduction, in Ballard, R.D., ed., Archaeological Oceanography: Princeton, New Jersey, Princeton University Press, p. ix–x. Berner, R.A., 1970, Sedimentary pyrite formation: American Journal of Science, v. 268, p. 1–23, doi: 10.2475/ajs.268.1.1. Björdal, C.G., and Nilsson, T., 2008, Reburial of shipwrecks in marine sediments: A long-term study on wood degradation: Journal of Archaeological Science, v. 35, p. 862–872, doi: 10.1016/j.jas.2007.06.005. Buxton, B., Ballard, R., Brennan, M., Coleman, D., Croff, K., Davis, D., Piechota, D., and Voronov, S., 2008, Byzantium beneath the Black Sea (poster): 2008 Meeting of the Archaeological Institute of America: Chicago, Archaeological Institute of America. Coleman, D.F., and Ballard, R.D., 2007, Submerged paleoshorelines in the southern and western Black Sea: Implications for inundated prehistoric archaeological sites, in Yanko-Hombach, V., Gilbert, A.S., Panin, N., and Dolukhanov, P.M., eds., The Black Sea Flood Question: Changes in Coastline, Climate and Human Settlement: Dordrecht, Netherlands, Springer, p. 671–696. Coleman, D.F., and Ballard, R.D., 2008, Oceanographic methods for underwater archaeological surveys, in Ballard, R.D., ed., Archaeological Oceanography: Princeton, New Jersey, Princeton University Press, p. 3–14.

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MANUSCRIPT ACCEPTED BY THE SOCIETY 22 JUNE 2010

Printed in the USA

The Geological Society of America Special Paper 473 2011

Pontic-Baltic pathways for invasive aquatic species: Geoarchaeological implications Ilya V. Buynevich* Geology & Geophysics Department, Woods Hole Oceanographic Institution, MS 22, Woods Hole, Massachusetts 02543, USA Aldona Damušytė Department of Quaternary Geology, Lithuanian Geological Survey, 35 S. Konarskio St., Vilnius, LT-03123, Lithuania Albertas Bitinas Sergej Olenin Coastal Research and Planning Institute, Klaipėda University, 84 H. Manto St., Klaipėda, LT-92294, Lithuania Jonas Mažeika Rimantas Petrošius Radioisotope Research Laboratory, Institute of Geology and Geography, 13 Ševčenkos St., Vilnius, LT-03223, Lithuania

ABSTRACT An accurate chronology for the exchange of aquatic species between water basins is important for paleoenvironmental reconstruction on both regional and continental scales. During the early Holocene, the range of zebra mussels, Dreissena polymorpha, was limited to the Black, Azov, Caspian, and Aral Seas, as well as the estuaries and lower and middle reaches of the Pontic-Caspian rivers. We present new findings that challenge the currently held view that this species migrated into the Baltic Sea watershed during the early 1800s through the canals joining the tributaries of rivers that drain into the Black and Baltic Sea basins. Geological investigations along the southeast Baltic Sea coast (Curonian and Vistula spits and lagoons) have uncovered shells of D. polymorpha that yielded radiocarbon ages older than 1000 radiocarbon yr B.P. We propose two scenarios to explain the new radiocarbon dates for D. polymorpha. The first scenario involves an anomalously large reservoir effect—as large as 600– 800 yr—however, several lines of evidence cast doubt upon the validity of such a large reservoir correction. The second scenario that might explain the old zebra mussel ages is the earlier arrival of Dreissena polymorpha into the Baltic region. Natural exchange may have been facilitated by the proximity of the tributaries draining the

*Present address: Department of Earth and Environmental Science, Temple University, 1901 N. 13th Street, Philadelphia, Pennsylvania 19122, USA; coast@ temple.edu. Buynevich, I.V., Damušytė, A., Bitinas, A., Olenin, S., Mažeika, J., and Petrošius, R., 2011, Pontic-Baltic pathways for invasive aquatic species: Geoarchaeological implications, in Buynevich, I.V., Yanko-Hombach, V., Gilbert, A., and Martin, R.E., eds., Geology and Geoarchaeology of the Black Sea Region: Beyond the Flood Hypothesis: Geological Society of America Special Paper 473, p. 189–196, doi: 10.1130/2011.2473(12). For permission to copy, contact [email protected]. © 2011 The Geological Society of America. All rights reserved.

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Buynevich et al. Pontic and Baltic watersheds. Human-mediated transport is also considered in association with Viking voyages from the Baltic to the Black and Caspian Seas between A.D. 800 and 1000, and the riverine trade exchange during the Lithuanian expansion into the Pontic steppe in subsequent centuries. It is likely that both scenarios played a role, with implications for late Holocene biogeography and paleoecology of the Pontic-Caspian and Baltic basins.

INTRODUCTION Throughout geologic history, the spread of aquatic species with broad environmental tolerances into foreign water basins has had profound consequences for the ecological balance between the native and newly immigrant species. In western Europe and North America, invasive species have caused adverse ecosystem changes, bringing with them profound societal consequences despite the fact that the timing and pathways of their transfer are not always known (Carlton, 1992; Vanderploeg et al., 2002; Daunys et al., 2006; Zaiko et al., 2008). Accurate dating of the appearance of marine, brackish, or freshwater species in various basins has important implications for paleogeography, paleoecology, and the chronology of enclosing sediments (Vitousek, 1990; Olenin and Leppäkoski, 1999; Behrends et al., 2005; Buynevich, 2007). There are also important archaeological components regarding inland navigation routes and possible human-mediated exchange of aquatic species between water basins (Petersen et al., 1992). In this study, we present new evidence of a potentially much earlier appearance of the zebra mussel, Dreissena polymorpha

Baltic Sea

3688 36884

Curonian Curon onian an Spit (Fig.. 2) (Fig.

Lithuania 26221 2622

Curonian Curon onian an Lagoon Lagoo

Nemunas River Rive ver delta elta d Baltic Sea

Vistula Spit

STUDY AREA

Russia

Poland

(Pallas, 1771), a native of the Black-Azov-Caspian-Aral Sea region, in the coastal waters of the Baltic Sea (Damušytė et al., 2007; Bitinas et al., 2008; Buynevich et al., 2010). D. polymorpha inhabits freshwater basins, but its tolerance of brackish water and its high reproduction rates have made it one the most widely known invaders of freshwater aquatic ecosystems (MacIsaac, 1996; Olenin and Leppäkoski, 1999). During the early Holocene, the range of D. polymorpha was limited to the Black, Azov, Caspian, and Aral Seas, as well as the estuaries and lower and middle reaches of the Pontic-Caspian rivers. The species was unintentionally introduced by humans into northwestern Russia, central and western Europe, Scandinavia, Britain, Ireland, and North America (Johnson and Carlton, 1996; Karatayev et al., 1997; Vanderploeg et al., 2002); however, the exact timing of its arrival in northern European watersheds is still in question. Our new findings challenge the currently held view that this species migrated into the Baltic watershed during the early 1800s through the canals joining the tributaries of rivers draining into the Black and Baltic Sea basins (Karatayev et al., 1997; Olenin et al., 1999). The aim of the present paper is to examine a new

50 km Black Sea

Figure 1. Map of the study area within Europe (inset) and satellite image of the southeastern Baltic Sea coast (Google Earth Image) showing the locations of sites where Dreissena polymorpha samples were collected. Sites include the bottom of the Curonian (Lithuania) and Vistula (Poland) Lagoons (triangles), outcrops of lagoonal sediments on the Curonian and Vistula Spits (circles), and two boreholes in the Curonian Lagoon (numbered boreholes).

Pontic-Baltic pathways for invasive aquatic species set of radiocarbon ages of Dreissena polymorpha from different environmental settings along the Baltic Sea coast and to discuss their geoarchaeological implications, including plausible migration pathways out of the Black Sea drainage basin. ECOLOGY Dreissena polymorpha is an ideal candidate for an invasive aquatic species due to its unusual ecology: (1) long-term survival during transport, and (2) rapid adaptation to a new environment, given a tolerable range of physico-chemical factors (Olenin et al., 1999; Vanderploeg et al., 2002; Daunys et al., 2006). The mollusk can tolerate salinity from fresh to 10–12‰, water temperatures up to 29 °C, and current velocities up to 2 m/s (Strayer and Smith, 1993). Dreissena has an unusual external fertilization style with a planktonic veliger larval stage, which allows for rapid dispersal (Jantz and Neumann, 1998). Once attached (sessile stage), zebra mussels not only thrive on outfall conduits and immersed parts of vessels, but they are reported to survive in dry docks on the hulls of commercial watercraft (Keevin et al., 1992). Therefore, the ecology of this mollusk species makes it one of the best suited for long-range transport, including a human-mediated vector through inland navigation (MacIsaac, 1996; Minchin et al., 2003). GEOLOGICAL SETTING Our study focused on several geological settings:(1) the bottom of the Curonian (Lithuania) and Vistula (Poland) Lagoons; (2) boreholes within the Curonian Lagoon; and (3) exposed

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lagoonal sediments behind the Parnidis Dune, Curonian Spit, in Lithuania, and the Vistula Spit, Kaliningrad region, in Russia and northern Poland (Fig. 1). The Curonian Lagoon is a shallow, nearly freshwater basin separated from the Baltic Sea by the Curonian Spit, a 97-km-long barrier spit extending in a SW-NE orientation. The spit possesses the highest coastal dunes in northern Europe (more than 60 m above sea level), and, in places, lagoonal sediments have been extruded above lagoon level due to the migration and loading of massive dunes (e.g., the Parnidis Dune site; Gudelis, 1998; Fig. 2). The Nemunas River enters the Curonian Lagoon and drains lowlands underlain by glacial deposits over Cretaceous bedrock (Bitinas et al., 2002; Guobytė, 2002). Dreissena shells were collected in situ from both the upper sandy layer of old lagoonal mud exposures (Fig. 3) and a section of lagoonal muds extruded due to dune loading in 2007 (Fig. 2; Buynevich et al., 2010). It is believed that D. polymorpha penetrated into the Baltic Sea basin (Curonian and Vistula Lagoons) only in the early 1800s via human-mediated invasion corridors: the Oginskiy Canal connecting the Nemunas and Pripet Rivers (opened in 1769) and the Bug-Pripet Canal (1775) (Karatayev et al., 1997; Olenin et al., 1999). METHODS Living and fossil Dreissena polymorpha shells from the southeastern Baltic Sea coast were photographed (Fig. 3) and catalogued. Well-preserved articulated and disarticulated shells were collected from the bottom of the Curonian and Vistula Lagoons, two boreholes in the Curonian Lagoon, as well as the

old exposure (Fig. 3)

Curonian Lagoon

Figure 2. Exposures of lagoonal strata at the base of the active Parnidis Dune along the landward margin of the Curonian Spit, Lithuania (looking southeast). Numerous in situ shells occur in both the old vegetated exposure and a recent (2007) outcrop. Samples dated from each exposure yield very similar radiocarbon ages. See Figure 1 for location.

new exposure

Parnidis Dune (active slipface)

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Buynevich et al. performed at the National Ocean Sciences AMS (NOSAMS) Laboratory at the Woods Hole Oceanographic Institution. This approach enabled us to compare the results of the two methods and assess the possibility of age contamination in bulk-sample dating due to pretreatment limitations of large samples. All 14 C ages were calibrated using OxCal, with a standard 400 yr reservoir correction applied (Table 1; Fig. 4). In the following sections, all ages are presented as years A.D. to facilitate comparison with historical events. RESULTS

Figure 3. Photograph of the older lagoonal outcrop showing large Dreissena polymorpha mollusks in situ; most are articulated (see Figs. 1 and 2 for location). The shells occur in the top sandy horizon and are absent from the lagoonal sediment below.

upper sections of the lagoonal marl outcrops behind the Parnidis Dune (Figs. 1 and 2). The shells ranged from 0.3 to 0.5 cm (juveniles in boreholes) to 3 cm in size (Fig. 3). Four samples were dated using a bulk method, a liquid scintillation counting (LSC) on bulk carbonate material of tens of D. polymorpha valves. Seven shells were dated using the accelerated mass spectrometry (AMS) method on a single valve (or several juvenile valves). The bulk dating on a collection of shells from the same lagoon bottom location or stratigraphic unit was performed at the Radioisotope Research Laboratory of the Lithuanian Institute of Geology and Geography. The AMS dating was

Radiocarbon dating of D. polymorpha valves from a variety of environmental settings forms the basis for our investigation into the chronology of the species along the Baltic Sea coast of Lithuania and the Russian Federation (Fig. 1). Valves of this species have also been found in late Holocene lagoonal deposits along the western coast of Poland (T. Radziejewska, 2008, personal commun.). The results of bulk and AMS age determinations based on the present study are summarized next (2σ OxCalcalibrated ages with 400 yr reservoir correction). Bulk-Sample Method Samples consisting of multiple shells of D. polymorpha collected from the bottom of the Curonian and Vistula Lagoons yielded bulk ages ranging from A.D. 220–945 (Table 1; Fig. 4). Four samples of other mollusks (Unio sp. and Viviparus sp.) from the lagoon floor also produced dates older than A.D. 900. A subfossil shell from the Parnidis Dune outcrop yielded a wide calibrated age range of A.D. 550–1100. A bulk sample from the Vistula Lagoon produced the oldest age of A.D. 30–410.

TABLE 1. RADIOCARBON DATES OF DREISSENA POLYMORPHA FROM SOUTHEAST BALTIC SEA COAST Lab no.

Geological context #

†

14

C age (yr, ±2σ error)

Calibrated age (cal yr A.D.)

Bulk method 150–700 VS-1565 Top of exposed mud, CL 1570 ± 120 VS-1642 Modern*, CL bottom 1490 ± 80 800–1090 VS-1681 Modern*, CL bottom 1550 ± 50 810–1020 VS-1703 Modern*, VL bottom 2130 ± 60 30–410 †† AMS method OS-57403 Top of exposed mud, CL 1250 ±30 1070–1270 OS-69634 Top of exposed mud, CL 1300 ± 30 1060–1180 OS-57405 Exposed mud, VL 1340 ± 30 1040–1130 OS-59536 Borehole 26221, CL 1600 ± 25 810–840 OS-57404 Borehole 26221, CL 285 ± 30 1890–modern OS-69635 Borehole 36884, CL 110 ± 25 1870–modern OS-69633 Live bivalve, CL bottom 885 ± 30 1430–1620 † CL—Curonian Lagoon, Lithuania; VL—Vistula Lagoon, Russia/Poland (outcrops of exposed/extruded lagoonal mud are now located above sea level). § Calendar age (cal yr A.D.), including 400 yr standard reservoir correction. # Radioisotope Research Laboratory, Institute of Geology and Geography, Lithuania. *Some bulk samples may include fossil forms exposed on lagoon bottom. †† National Ocean Sciences Accelerator Mass Spectrometry (NOSAMS) facility, Woods Hole Oceanographic Institution, USA.

§

Pontic-Baltic pathways for invasive aquatic species AMS Method In contrast to the bulk dates, the AMS samples yielded younger ages and considerably smaller errors (Table 1; Fig. 4). One valve of a living Dreissena with byssal threads was dated to A.D. 1430–1620. Single valves of in situ articulated D. polymorpha from the old and the most recent lagoonal mud outcrops at Parnidis Dune (Figs. 2 and 3) were dated to A.D. 1070–1270 and A.D. 1060–1180, respectively. The Vistula Spit sample produced a similar age of A.D. 1040–1130. Two samples from a depth of 6.7 m in borehole 26221 were dated to A.D. 810–950 and A.D. 1890– present day, and several valves of juveniles sampled from a depth of 5.5 m in borehole 36884 yielded a nearly modern age (Fig. 4). DISCUSSION The results of radiocarbon dating indicate that a number of determined ages for Dreissena polymorpha are substantially older than the early 1800s. This period is currently regarded as a

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time of migration of the mussels through the canals, which were excavated to join the headwaters and tributaries of rivers flowing into the Black Sea (Dnieper and Bug Rivers) to those entering the bays and lagoons of the Baltic Sea (Nemunas and Daugava Rivers; Fig. 5). The new set of dates presented here calls for a revised chronology and an explanation for the older ages and wider range of dates. Local Reservoir Effect Scenario One of the considerations that may explain the relatively old ages of modern Dreissena polymorpha shells is an abnormally high reservoir effect of 600–800 yr (Table 1; Fig. 4). This suggests a local reservoir correction (ΔR) of 200–400 yr, in addition to the standard 400 yr correction. Living mollusks from inland lakes dated by one of the present authors (Mažeika) provide modern ages, so some factors unique to the Curonian Lagoon would have to be responsible for the large reservoir effect. Carbonate enrichment in surrounding sediments or input from the Nemunas

2400 2200

shells from lagoonal exposures (Curonian and Vistula Spits)

2000

Radiocarbon age (14C years)

1800 1600 1400 1200 1000 800

documented species introduction through canals

600 400 200

Viking voyages Lithuanian southerly expansion

Little Ice Age

0 BC/AD 0

200

400

600

800

1000

1200

1400

1600

1800

2000

Calibrated age (years AD) Figure 4. Plot of raw radiocarbon dates (14C yr) versus OxCal-calibrated ages (calendar yr A.D.) with a standard 400 yr reservoir correction (2σ error). Bulk sample results (Institute of Geology and Geography, Lithuania) are shown in gray; accelerated mass spectrometry (AMS) analyses (NOSAMS, Woods Hole Oceanographic Institution, USA) are shown in black. Note the younger ages and a substantially smaller error for AMS dates. The ages of key historical events are shown along the x-axis.

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River are still under consideration; however, this does not explain the old ages obtained on specimens from Vistula Spit exposures farther south (Fig. 1). In a recent study of lagoonal deposits along the Vistula Spit (Fig. 1), Bitinas et al. (2008) report a similarity between the bulk ages of lagoonal sediments (580–200 B.C.) and the embedded Unio pictorum fragments if a standard 300–400 yr reservoir correction was applied to the latter (560–380 B.C.). More evidence arguing against the abnormally large reservoir effect involves the young 14C dates of shells in boreholes, where they were likely forced from the surface sediments by the coring procedure. In addition, if the species arrived in the region 200 yr ago, a 400-to-700 yr range in age cannot be accounted for by a reservoir correction alone. The isotopic fractionation (δ13C) value of the seven AMS-dated shells ranges from = −9.97‰ to −5.20‰, which suggests a riverine/lagoonal environment (Keith and Weber, 1964; Lanting and van der Plicht, 1998) and implies relatively minor impact from a fully marine reservoir effect. Finally, the fact that several samples were dated to near modern ages (Fig. 4) argues against the regional reservoir correction. While the reservoir effect on shell age cannot be ruled out to par-

A

tially explain the old ages, an alternative explanation has to be considered that involves the earlier arrival of Dreissena polymorpha than the currently accepted time frame of the early 1800s. There are examples of reservoir (“hard-water”) effects from other sites in the Baltic region. Experimental analyses by Fischer and Heinemeier (2003) revealed a 1000-yr-old age for presentday freshwater mussels, as well as a 100–500 yr discrepancy between radiocarbon ages of food residue (fish and mollusks) on pottery from Denmark and the known archaeological dating. The effect of the cooking process and vessel composition on the 14 C age of the shells may also play a role in such cases, whereas only natural phenomena must be considered when explaining the abnormally old dates and wide age range of D. polymorpha in the present study. The influence of Cretaceous bedrock on the isotopic composition of groundwater and surface water of the Nemunas River must be explored further; however, it is still difficult to explain a number of near-modern shell dates from the Curonian Lagoon (Table 1; Fig. 4). This reservoir-effect scenario requires further confirmation, with potential comparisons to the results of other dating techniques (e.g., U-Th dating).

Viking territories AD 800-1100

B 820

Baltic Sea

854 880 STUDY AREA

Nemunas R.

Viking voyages

Dnieper R.

882

Pripet R.

Dniester R. Bug R.

Lithuanian territories (ca. 1420s)

941 839 Black Sea

854 – Viking voyage date (years AD)

Figure 5. Map of Eastern Europe showing the possible pathways for movement of aquatic species. (A) Circles outline regions where tributaries of the rivers to the Pontic and Baltic drainage basins are in natural proximity to each other (A.D. 1000 map). (B) Viking territories and major river and sea voyages during the interval A.D. 800–1000 (modified after Graham-Campbell, 2001) and the extent of Lithuanian territories during A.D. 1200–1500 (dashed outline; after Rowell, 1994; Gudavičius, 1999).

Pontic-Baltic pathways for invasive aquatic species Earlier Arrival Scenario One of the pathways for migration of D. polymorpha between the tributaries of the rivers draining into the Black Sea and those flowing into the Baltic Sea may have been a natural migration through poorly drained waterways. In several areas, the distance between tributary headwaters does not exceed 10– 15 km, and this proximity may have played a role in facilitating the migration, particularly during flood periods (Fig. 5A). An alternative scenario includes human-mediated pathways, such as transfer of D. polymorpha larvae or juveniles in waterstorage containers or adult mollusks attached to the hulls and anchors of small watercraft traveling from the Pontic-Caspian to the Baltic watersheds. As mentioned already, the ecology of young and adult mussels and their rapid recent expansion attest to the flexibility and subsequent adaptability of Dreissena as an invasive species (Johnson and Carlton, 1996; MacIsaac, 1996; Vanderploeg et al., 2002; Minchin et al., 2003). There is ample evidence of Viking river voyages via the Dnieper River and Black Sea to Constantinople (e.g., Eastern Line, Varanges to Greeks) during A.D. 820–941 (Figs. 4 and 5B), via the Volga River and Caspian Sea to Baghdad, and probably more local exchanges at an earlier time (Christiansen, 1997; Wise, 2005; Chartrand et al., 2006). The return voyages to the Baltic Sea may have afforded a potential pathway for the mussels. At a later time, there may have been other opportunities for transfer of Dreissena polymorpha and other species, primarily as a result of overland exchange of watercraft between countries occupying the headwater region. For example, the Grand Duchy of Lithuania and Lithuanian Allies and Dependencies extended from the Baltic Sea south to include the headwaters of the Dnieper River by A.D. 1263, ultimately reaching the Pontic steppe by A.D. 1420 (Figs. 4 and 5B; Rowell, 1994; Gudavičius, 1999). During this period, the riverine and overland pathways may have been one of the routes for zebra mussel larvae or attached mollusks. The human-mediated pathways would have resulted in a relatively rapid transport of the mussels, thus giving them an opportunity to adjust to new, but ecologically satisfactory, conditions in the Baltic lagoons. It is worth noting that only one sample dates to the period of A.D. 1300–1800, even after application of a standard 400 yr reservoir correction (Fig. 4). This interval closely corresponds to the Little Ice Age (A.D. 1350–1850) and may explain the general absence of D. polymorpha if the water temperature was cooler than optimal conditions. It is not clear why few samples from the Baltic coast date to this time. Given an earlier arrival scenario, the migration of the mussels through the canals in the late 1700s– early 1800s would have reintroduced them into the region at the end of the Little Ice Age. CONCLUSIONS Our new findings challenge the currently held view that Dreissena polymorpha migrated into the Baltic watershed dur-

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ing early 1800s through the canals joining the tributaries of rivers draining into the two basins. Geological investigations along the Curonian and Vistula Spits, as well as the Curonian Lagoon, provide a data set of D. polymorpha shells with radiocarbon ages greater than 1000 radiocarbon yr B.P. These new findings are substantially older than the previously documented introduction the of zebra mussel into the Baltic region (Olenin et al., 1999). We propose two scenarios to explain the new radiocarbon dates of D. polymorpha. The first scenario involves an anomalously large reservoir effect. Bulk radiocarbon ages of several present-day mollusks from the Curonian Lagoon range from A.D. 220 to 1525. This may be due to the fact that the lagoon water is enriched with old carbonates at a higher level than in the Baltic Sea and yields an age difference as large as 600– 800 yr. D. polymorpha from outcrops along the Curonian and Vistula Spits yielded AMS dates within the A.D. 640–870 range, suggesting a similar reservoir effect for the two lagoons. However, nearly modern ages of several other Dreissena samples from the Curonian Lagoon, a 400–700 yr spread in ages, and similarity between shell ages in different lagoons provide some basis for questioning the validity of an abnormally large reservoir effect scenario. While still considering the contribution of a local reservoir effect to both bulk and AMS ages, we propose a second scenario to explain the old ages: the possibility of an earlier arrival of Dreissena polymorpha into the Baltic region (Fig. 5). The natural exchange may have taken place during periods of floods or tributary capture facilitated by the proximity between rivers draining the Pontic and Baltic watersheds. Also, a human-mediated transport between the rivers may have occurred, including both Viking travels from the Baltic to the Black and Caspian Seas during A.D. 800–1000 and the riverine trade exchange during the Lithuanian expansion into the Pontic steppe in subsequent centuries. Regardless of which scenario proves to be correct, the present study has important geoarchaeological implications. Until the question is fully resolved, caution must be exercised by geologists and archaeologists before assigning a young age to deposits or cultural materials containing Dreissena polymorpha. Both the reservoir effect and earlier arrival scenarios should be considered when reconstructing the chronology of aquatic species exchange, with implications for late Holocene biogeography and paleoecology of the Pontic-Caspian and Baltic basins. ACKNOWLEDGMENTS This research was partially funded by the Ocean and Climate Change Institute of the Woods Hole Oceanographic Institution, The National Geographic Society, and the Lithuanian Geological Survey. We thank Donatas Pupienis and Anton Symonovich for their support in the field and numerous colleagues for productive discussions on the subject. D. Basso, K. Petersen, and A. Gilbert provided many useful comments on the manuscript. This paper is a contribution to the International Geological Correlation Programme Project 521.

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