Monsoon Evolution and Tectonics-Climate Linkage in East Asia (Geological Society Special Publication 342)

Monsoon Evolution and Tectonics–Climate Linkage in Asia

The Geological Society of London Books Editorial Committee Chief Editor

BOB PANKHURST (UK) Society Books Editors

JOHN GREGORY (UK) JIM GRIFFITHS (UK) JOHN HOWE (UK) RICK LAW (USA) PHIL LEAT (UK) NICK ROBINS (UK) RANDELL STEPHENSON (UK) Society Books Advisors

MIKE BROWN (USA) ERIC BUFFETAUT (FRANCE ) JONATHAN CRAIG (ITALY ) RETO GIERE´ (GERMANY ) TOM MC CANN (GERMANY ) DOUG STEAD (CANADA ) MAARTEN DE WIT (SOUTH AFRICA )

IUGS/GSL publishing agreement This volume is published under an agreement between the International Union of Geological Sciences and the Geological Society of London and arises from [IUGS commission/IGCP programme name & no here]. GSL is the publisher of choice for books related to IUGS activities, and the IUGS receives a royalty for all books published under this agreement. Books published under this agreement are subject to the Society’s standard rigorous proposal and manuscript review procedures.

It is recommended that reference to all or part of this book should be made in one of the following ways: CLIFT , P. D., TADA , R. & ZHENG , H. (eds) 2010. Monsoon Evolution and Tectonics –Climate Linkage in Asia. Geological Society, London, Special Publications, 342. WAN , S., CLIFT , P. D., LI , A., LI , T. & YIN , X. 2010. Geochemical records in the South China Sea: implications for East Asian summer monsoon evolution over the last 20 Ma. In: CLIFT , P. D., TADA , R. & ZHENG , H. (eds) Monsoon Evolution and Tectonics–Climate Linkage in Asia. Geological Society, London, Special Publications, 342, 245 –263.

GEOLOGICAL SOCIETY SPECIAL PUBLICATION NO. 342

Monsoon Evolution and Tectonics – Climate Linkage in Asia

EDITED BY

P. D. CLIFT University of Aberdeen, UK

R. TADA University of Tokyo, Japan

and H. ZHENG University of Nanjing, China

2010 Published by The Geological Society London

THE GEOLOGICAL SOCIETY The Geological Society of London (GSL) was founded in 1807. It is the oldest national geological society in the world and the largest in Europe. It was incorporated under Royal Charter in 1825 and is Registered Charity 210161. The Society is the UK national learned and professional society for geology with a worldwide Fellowship (FGS) of over 9000. The Society has the power to confer Chartered status on suitably qualified Fellows, and about 2000 of the Fellowship carry the title (CGeol). Chartered Geologists may also obtain the equivalent European title, European Geologist (EurGeol). One fifth of the Society’s fellowship resides outside the UK. To find out more about the Society, log on to www.geolsoc.org.uk. The Geological Society Publishing House (Bath, UK) produces the Society’s international journals and books, and acts as European distributor for selected publications of the American Association of Petroleum Geologists (AAPG), the Indonesian Petroleum Association (IPA), the Geological Society of America (GSA), the Society for Sedimentary Geology (SEPM) and the Geologists’ Association (GA). Joint marketing agreements ensure that GSL Fellows may purchase these societies’ publications at a discount. The Society’s online bookshop (accessible from www.geolsoc.org.uk) offers secure book purchasing with your credit or debit card. To find out about joining the Society and benefiting from substantial discounts on publications of GSL and other societies worldwide, consult www.geolsoc.org.uk, or contact the Fellowship Department at: The Geological Society, Burlington House, Piccadilly, London W1J 0BG: Tel. þ44 (0)20 7434 9944; Fax þ44 (0)20 7439 8975; E-mail: [email protected]. For information about the Society’s meetings, consult Events on www.geolsoc.org.uk. To find out more about the Society’s Corporate Affiliates Scheme, write to [email protected]. Published by The Geological Society from: The Geological Society Publishing House, Unit 7, Brassmill Enterprise Centre, Brassmill Lane, Bath BA1 3JN, UK (Orders: Tel. þ44 (0)1225 445046, Fax þ44 (0)1225 442836) Online bookshop: www.geolsoc.org.uk/bookshop The publishers make no representation, express or implied, with regard to the accuracy of the information contained in this book and cannot accept any legal responsibility for any errors or omissions that may be made. # The Geological Society of London 2010. All rights reserved. No reproduction, copy or transmission of this publication may be made without written permission. No paragraph of this publication may be reproduced, copied or transmitted save with the provisions of The Copyright Licensing Agency Ltd, Saffron House, 6 –10 Kirby Street, London EC1N 8TS, UK. Users registered with the Copyright Clearance Center, 222 Rosewood Drive, Danvers, MA 01923, USA: the item-fee code for this publication is 0305-8719/10/$15.00. British Library Cataloguing in Publication Data A catalogue record for this book is available from the British Library. ISBN 1-86239-310-3 Typeset by Techset Composition Ltd, Salisbury, UK Printed by CPI Antony Rowe Ltd, Chippenham, UK Distributors North America For trade and institutional orders: The Geological Society, c/o AIDC, 82 Winter Sport Lane, Williston, VT 05495, USA Orders: Tel. þ1 800-972-9892 Fax þ1 802-864-7626 E-mail: [email protected] For individual and corporate orders: AAPG Bookstore, PO Box 979, Tulsa, OK 74101-0979, USA Orders: Tel. þ1 918-584-2555 Fax þ1 918-560-2652 E-mail: [email protected] Website: http://bookstore.aapg.org India Affiliated East-West Press Private Ltd, Marketing Division, G-1/16 Ansari Road, Darya Ganj, New Delhi 110 002, India Orders: Tel. þ91 11 2327-9113/2326-4180 Fax þ91 11 2326-0538 E-mail: [email protected]

Contents C LIFT , P. D., T ADA , R. & Z HENG , H. Monsoon evolution and tectonics– climate linkage in Asia: an introduction

1

C HEN , M.-T., C HANG , Y.-P., Y U , P.-S. & S HIAU , L.-J. Marine records of East Asian monsoon variability over the past 5 Ma

5

D E , S., S ARKAR , S. & G UPTA , A. K. Orbital and suborbital variability in the equatorial Indian Ocean as recorded in sediments of the Maldives Ridge (ODP Hole 716A) during the past 444 ka

17

L U , H., W ANG , X. & L I , L. Aeolian sediment evidence that global cooling has driven late Cenozoic stepwise aridification in central Asia

29

T ADA , R., Z HENG , H., S UGIURA , N., I SOZAKI , Y., H ASEGAWA , H., S UN , Y., Y ANG , W., W ANG , K. & T OYODA , S. Desertification and dust emission history of the Tarim Basin and its relation to the uplift of northern Tibet

45

Z HENG , H., T ADA , R., J IA , J., L AWRENCE , C. & W ANG , K. Cenozoic sediments in the southern Tarim Basin: implications for the uplift of northern Tibet and evolution of the Taklimakan Desert

67

S UN , Y., A N , Z. & C LEMENS , S. C. Non-stationary response of Plio-Pleistocene East Asian winter monsoon variation to ice volume forcing

79

S TEVENS , T. & L U , H. Radiometric dating of the late Quaternary summer monsoon on the Loess Plateau, China

87

S AKAI , T., S ANEYOSHI , M., T ANAKA , S., S AWADA , Y., N AKATSUKASA , M., M BUA , E. & I SHIDA , H. Climate shift recorded at around 10 Ma in Miocene succession of Samburu Hills, northern Kenya Rift, and its significance

109

H UH , Y. Estimation of atmospheric CO2 uptake by silicate weathering in the Himalayas and the Tibetan Plateau: a review of existing fluvial geochemical data

129

S ANYAL , P. & S INHA , R. Evolution of the Indian summer monsoon: synthesis of continental records

153

C LIFT , P. D., G IOSAN , L., C ARTER , A., G ARZANTI , E., G ALY , V., T ABREZ , A. R., P RINGLE , M., C AMPBELL , I. H., F RANCE -L ANORD , C., B LUSZTAJN , J., A LLEN , C., A LIZAI , A., L U¨ CKGE , A., D ANISH , M. & R ABBANI , M. M. Monsoon control over erosion patterns in the Western Himalaya: possible feed-back into the tectonic evolution

185

H OANG , L. V., C LIFT , P. D., S CHWAB , A. M., H UUSE , M., N GUYEN , D. A. & Z HEN , S. Large-scale erosional response of SE Asia to monsoon evolution reconstructed from sedimentary records of the Song Hong-Yinggehai and Qiongdongnan basins, South China Sea

219

W AN , S., C LIFT , P. D., L I , A., L I , T. & Y IN , X. Geochemical records in the South China Sea: implications for East Asian summer monsoon evolution over the last 20 Ma

245

M OTOI , T. & C HAN , W.-L. Colder Subarctic Pacific with larger sea ice caused by closure of the Central American Seaway and its influence on the East Asian monsoon: a climate model study

265

L UNT , D. J., F LECKER , R. & C LIFT , P. D. The impacts of Tibetan uplift on palaeoclimate proxies

279

K ITOH , A., M OTOI , T. & A RAKAWA , O. Climate modelling study on mountain uplift and Asian monsoon evolution

293

Index

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Monsoon evolution and tectonics – climate linkage in Asia: an introduction PETER D. CLIFT1*, RYUJI TADA2 & HONGBO ZHENG3 1

School of Geosciences, Meston Building, University of Aberdeen, Aberdeen, AB24 3UE, UK 2

Department of Earth and Planetary Science, University of Tokyo, Science Building #1, 7-3-1 Hongo, Tokyo 113-0033, Japan

3

Institute of Surficial Geochemistry, School of Earth Science and Engineering, Nanjing University, 22 Hankou Road, Nanjing, People’s Republic of China *Corresponding author (e-mail: [email protected])

Interactions between the solid Earth and climate, both on local and global scales are increasingly being considered as important within the sphere of the Earth and ocean sciences. For example, it has long been recognized that opening and closure of oceanic gateways, as a result of continental breakup and collision processes, can lead to changes in oceanic circulation patterns and so to changes in climate (Kennett 1977; Haug et al. 2001; von der Heydt & Dijkstra 2006). In addition, uplift of mountain chains can disrupt atmospheric circulation by deflecting the jet stream and altering planetary climatic belts (Tada 2004), as well as generating orographic rainfall concentration and rain shadows in the immediate vicinity of mountainous topography (Jiang et al. 2003). However, the most dramatic example of the solid Earth affecting climate is the proposed relationship between the growth of the topography in Central Asia during the Cenozoic and the intensification of the Asian monsoon. Asia is not the only continent to have a monsoon, but this monsoon is by far the most powerful and is driven by the temperature differences between the Eurasian continent and the Indian and Pacific Oceans (Webster et al. 1998; Clift & Plumb 2008), which causes a circulation reversal to the normal Hadley circulation in South and East Asia during the summer. In particular, growth of the Tibetan Plateau has been cited as being a trigger for a much stronger summer monsoon than might otherwise be predicted (Prell & Kutzbach 1992; Molnar et al. 1993). Numerical models suggest that growth of the plateau, beyond around half its present height would be important in greatly increasing rainfall along the central parts of the Himalayan mountain front (Kitoh 2004). In turn the greater precipitation affects the erosion of the mountain front and this may affect the tectonic evolution of the mountain through increasing and focusing exhumation, thus forming a positive feedback loop. Because erosion is required to cause exhumation in compressional settings rainfall

patterns can be important in controlling the overall structure of the ranges (Willett 1999). Indeed, areas of rapid exhumation in the Himalaya have been correlated with zones of the most intense summer rains (Thiede et al. 2004). Erosion not only affects the mountains themselves, but delivers large volumes of sediment to the ocean. The faster erosion of rock under the influence of intensified rain may result in enhanced chemical weathering under a humid climate in the flood plains of the foreland. These chemical reactions may be responsible for the consumption of atmospheric CO2, a greenhouse gas, which in turn can drive long-term global cooling (Raymo & Ruddiman 1992), thus forming a negative feedback loop. Furthermore, burial of organic carbon in deep-sea submarine fans may also cause a reduction in CO2 through removing large volumes of carbon from the normal carbon recycling process (Raymo 1994; Galy et al. 2007). Clearly interactions between the solid Earth, ocean and atmosphere are more complex than originally envisaged. As a result, in 2003 IGCP initiated Project 476 in order to examine the interactions between the tectonic evolution of Asia and the Asian monsoon system, as well as assessing their impact on the palaeoceanography of the Asian marginal seas. The IGCP group met five times, with the initial results published in a special collection of papers in volumes 241 and 247 of Palaeogeography Palaeoclimatology Palaeoecology in 2006 and 2007. In this volume we have collected the final results of our investigations. We thank UNESCO for their sponsorship of this initiative and the meeting hosts in Tokyo, Vladivostok, Busan and Shanghai where the group had its annual meetings. Marine oceanographic records were some of the first data sources used to date the onset and measure the development of the Asian monsoon and in this collection we present a number of new studies that provide additional details on how and why the monsoon has varied in the past over a number of

From: Clift, P. D., Tada, R. & Zheng, H. (eds) Monsoon Evolution and Tectonics –Climate Linkage in Asia. Geological Society, London, Special Publications, 342, 1–4. DOI: 10.1144/SP342.1 0305-8719/10/$15.00 # The Geological Society of London 2010.

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timescales. Chen et al. review proxies for marine sea surface temperature, sea surface salinity, and organic productivity from the South China and East China Seas in order to determine how the monsoon has changed in East Asia since 5 Ma. These authors show that the timing of monsoon variations based on sea surface temperature reconstructions have proxy-dependent discrepancies, which need to be resolved if we are to derive a more robust monsoon history. The monsoon is dominated by millennial-scale hydrographic and productivity variations, with warm, less saline water indicative of a stronger summer monsoon during interstadials, and colder, more saline waters in stadials. Working on even shorter time scales De et al. examined a 450 ka record from the region of the Maldives in the Arabian Sea, western Indian Ocean. This study employed high resolution proxy records of planktic foraminifera and pteropods from Ocean Drilling Program (ODP) Hole 716A. They revealed large-scale changes in summer monsoon wind intensity during the period 450– 300 ka, followed by a marked weakening after 300 ka. Since that time the monsoon seems to have varied on orbital and suborbital timescales, but generally has been less variable than before 300 ka. This collection brings together a number of contributions from terrestrial, mostly windblown records for continental climate in Central Asia, building on some of the earlier work from the Chinese Loess Plateau that showed strong winter monsoon intensification after 8 Ma (An et al. 2001; Guo et al. 2004; Zheng et al. 2004). Lu et al. present a new data compilation of aeolian records spanning the Early Miocene to late Pleistocene in central China, which they use to determine whether drying of the mid-continent was caused by global cooling or topographic uplift of the Tibetan Plateau, the latter causing a rain shadow effect. They demonstrate a close association of drying proxies with global cooling and suggest a reduced role for topographic growth in climate change. Instead these authors envisage that a reorganization of atmospheric circulation and ocean currents may play a vital role in the cooling driving continental desiccation. An alternative view of this process is provided by Tada et al. who examined a section of Neogene fluvial sedimentary rock on the southwestern margin of the Tarim Basin where deposition of aeolian siltstone started at 4.6 Ma. Tada et al. demonstrate that deposition of aeolian siltstone became significant after 4.0 Ma when tilting of the strata started and deposition of conglomerate became significant at the foot of the western Kunlun Mountains. They show a close linkage between tectonic uplift, erosion, sediment accumulation on alluvial fans, and dust emission from the surface of alluvial fans on the northwestern margin of Tibet that requires a linkage between uplift and

desertification. In a related study, Zheng et al. examined the Miocene–Recent sediments along the northern edge of the Tibetan Plateau in the Tarim Basin. Sections dated as 8 Ma at the base grade upwards from finer and more distal clastic sediments into coarse alluvial fan deposits related to the uplift of northern Tibet. At the same time aeolian dunes and playa lakes started to form after 8 Ma, suggesting an enclosed desert basin from that time, suggesting that plateau uplift may be causing the regional climatic drying. Zheng et al. conclude that the Taklimakan Desert and the regional climate regime may have been fully developed by the Early Pliocene. Shorter-term controls on winter monsoon intensity were investigated by Sun et al. who used Chinese loess and deep-sea sediments to address the issue of how the winter monsoon became coupled to global ice volume change. Comparison of the two data sets shows that over long timescales the winter monsoon exhibited strong coupling with global ice volume after 2.1 Ma, while at orbital timescales the winter monsoon variations started to be influenced by global ice volume change at c. 3.3 Ma. None the less, the situation is clearly more complicated because mismatches between these two records over both shorter and longer timescales require that additional processes played a role, and not just ice volume alone. In a linked study, Stevens & Lu examined the basis of using the loess as a monsoon proxy record. They used improved optically stimulated luminescence (OSL) age control to argue that use of CaCO3 as a proxy in the loess is complicated by the multiple influences on its formation. However, changes in magnetic susceptibility can be used as a proxy for summer monsoon induced pedogenesis. Using this information these authors demonstrate decreasing summer East Asian monsoon strength from 50 to 18 ka. However, between 9 and 6 ka, magnetic susceptibility increased abruptly and dramatically at the sites, indicative of a strong summer monsoon, albeit a little later than is generally recognized elsewhere. These findings suggest that there is a close correspondence with ice volume rather than solar insolation. Sakai et al. provide constraints on the co-evolution of the African Monsoon during the Miocene. These authors document a change in the East African (Kenyan) environment from a red soil-dominated condition to lacustrine and deltaic facies with open woodland/savannah mammalian fauna. This change is interpreted to reveal a shift from a dry climate with seasonal precipitation to a climate with strong seasonality, coincident with documented changes in the Asian monsoon. This study suggests that the change could be linked to the spread of a stronger Indian summer monsoon into Africa, but also proposes that the

INTRODUCTION

enhanced moisture transport could be simply linked to movement of the Intertropical Convergence Zone (ITCZ), synchronized with the Indian summer monsoon. Several studies consider the impact that changing monsoon strength has on the processes of continental erosion and chemical weathering. Huh used chemical data from Himalayan and Tibetan rivers to conclude that carbonate weathering dominates the major element composition of these rivers. She shows that there is significant variability in the silicate contributions, which are highest in the rivers of the Himalayan syntaxes and in the Yamuna, Alaknanda-Bhaghirathi and Kosi tributaries of the Ganges. The Gangetic tributaries seem to supply uniquely radiogenic 87Sr to the ocean, but the silicate weathering rates are comparable to other major rivers draining orogenic zones. This requires that 87Sr/86Sr ratios are decoupled from silicate weathering rates. Sanyal & Sinha examined the impact that monsoonal variations have had on erosion and sedimentation in the Himalayan foreland basin. These authors used a series of sedimentary proxies to propose that the monsoon intensified in a series of stages, with peaks at 10.5, 5.5 and 3 Ma after which the summer monsoon decreased to the modern strength. More recently the Ganges flood plain appears to have aggraded periodically between 27 and 90 ka, but then degraded under the influence of weaker monsoonal precipitation around the Last Glacial Maximum. Sanyal & Sinha propose that as the monsoon strengthened between 15 and 5 ka the flood plains experienced incision and badland formation up until the time of maximum drying around 3 ka. Further west the erosional response to monsoon intensification since the Last Glacial Maximum is documented by Clift et al. In this study a series of thermochronometers were applied to sediments cored in the Indus delta to show that a strengthening monsoon in the Early Holocene, after 11 ka, caused faster erosion and a shift in the location of that erosion from the Karakoram and Transhimalaya to the wetter, more southerly Lesser Himalaya. These authors speculate that if this climatic control on erosion extends over longer periods of geological time then this process could be the principle trigger for exhumation in the Greater Himalaya, rather than plate tectonic forces. Contrasting with this study is the evidence from Hoang et al. in the Gulf of Tonkin in the South China Sea. Their study was based on a seismic stratigraphic analysis of sediments in the Song Hong-Yinggehai Basin formed as a pull-apart basin at the southern end of the continental Red River Fault Zone. The basin is filled by sediments delivered by the Red River, allowing periods of faster erosion after 3–4 Ma and between 10 and 15 Ma to be identified and linked to times of strong summer monsoon.

3

However, the earliest pulse of fast erosion at 29.5 –21 Ma may have been triggered by tectonic rock uplift along the Red River Fault Zone. Hoang et al. show that chemical weathering has gradually decreased in SE Asia after c. 25 Ma, probably because of decreasing global temperatures, whereas physical erosion became stronger, especially after c. 12 Ma. Wan et al. attempt to reconstruct past changes in East Asian summer and winter monsoons during the past 20 Ma based on major and minor element chemistry of the marine sediments from the northern South China Sea. They demonstrated a general decreasing trend of summer monsoon strength since 14 Ma and an increasing trend of the winter monsoon since 20 Ma. These authors argued that evolution of the East Asian summer and winter monsoons are decoupled, and that evolution of the East Asian summer monsoon is more strongly controlled by global cooling rather than tectonically driven plateau uplift. Climate modelling now forms an important dimension to any attempt to understand the Asian monsoon. Models linking monsoon intensification to Tibetan Plateau growth are well recognized; however, little detail is known about the sensitivity of the monsoon to the uplift and about the effect of other factors, such as the retreat of shallow seas from Central Asia (Ramstein et al. 1997). As part of this collection Motoi et al. examine the effect that the Central American Seaway has had on the East Asian monsoon by considering the climatic state predicted for the Panama seaway being closed, open, and then re-closed. Their models predict formation of a permanent halocline in the Sub-Arctic Pacific as a result of the termination of saline water transport through the seaway, leading to extensive sea ice in winter and colder summer sea surface temperatures. These authors propose that the East Asian monsoon would be weakened in winter and strengthened in summer as a result of closing the Central American Seaway. In contrast Lunt et al. directly address the impacts that Tibetan surface uplift might be expected to have on climate in Asia and predict which proxies measured in oceanic sediments might be most affected by the progressive increase. Although sea surface temperatures in the western Pacific, South China Sea and Indian Ocean are generally insensitive to Tibetan uplift, vegetation in the region of the plateau itself is affected, as is discharge from the Yangtze, Pearl, and especially the Ganges/Brahmaputra Rivers. Kitoh et al. also examined the impact of plateau uplift using a general circulation model (GCM). Their model suggests that when there is no plateau present monsoon precipitation is confined in the deep tropics during northern hemispheric summer. However, as the Tibetan Plateau is uplifted, rainfall increases and falls instead inland over the southeastern Tibetan Plateau. Like Lunt et al. these models

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predict more fluvial discharge and thus a significant decrease of sea surface salinities over the Bay of Bengal, the South and East China Seas, as well as the Yellow Sea. These modelling studies reinforce the belief that integrating observational studies and models leads to improvements in both spheres, in particular leading to more targeted and better focused proxy studies. This IGCP has advanced the studies of climate– tectonic interactions in the case of the Asian monsoon but there is still much work to be done. The initial intensification of the monsoon has yet to be defined, and debate continues about which proxies are the best suited for determining this. Better, higher-resolution climate models are needed that treat the plateau not as a homogenous block, but as the laterally changing and growing edifice we believe it to be (Tapponnier et al. 2001; Royden et al. 2008). Advances also require more complete sedimentary erosional records, which are handicapped in South Asia by the lack of an Oligocene record throughout much of the Himalayan foreland (Najman 2006). This latter goal will require deep, scientific drilling of the Indian Ocean submarine fans, together with the deltas of SE and East Asia. Integration of oceanic records, atmospheric sciences and tectonic data from onshore lies at the heart of this work. Some of these goals are now being addressed by a new IGCP, 581, which addresses the evolution of Asian river systems and considers both how these are influenced by climatic and tectonic processes, as well as helping us understand these processes through study of the sediments that they have and continue to carry to the Asian marginal seas.

References An, Z., Kutzbach, J. E., Prell, W. L. & Porter, S. C. 2001. Evolution of Asian monsoons and phased uplift of the Himalaya-Tibetan Plateau since late Miocene times. Nature (London), 411, 62– 66. Clift, P. D. & Plumb, R. A. 2008. The Asian Monsoon: Causes, History and Effects. Cambridge University Press, Cambridge. Galy, V., France-Lanord, C., Beyssac, O., Faure, P., Kudrass, H.-R. & Palhol, F. 2007. Efficient organic carbon burial in the Bengal fan sustained by the Himalayan erosional system. Nature, 450, 407– 411, doi: 10.1038/nature06273. Guo, Z., Peng, S., Hao, Q., Biscaye, P. E., An, Z. & Liu, T. 2004. Late Miocene– Pliocene development of Asian aridification as recorded in the Red-Earth Formation in northern China. Global and Planetary Change, 41, 135 –145. Haug, G. H., Tiedemann, R., Zahn, R. & Ravelo, A. C. 2001. Role of Panama uplift on oceanic freshwater balance. Geology, 29, 207–210. Jiang, J. H., Wu, D. L., Eckermann, S. D. & Ma, J. 2003. Mountain waves in the middle atmosphere: microwave

limb sounder observations and analyses. Advances in Space Research, 32, 801–806. Kennett, J. P. 1977. Cenozoic evolution of Antarctic glaciation, the Circum-Antarctic Ocean, and their impact on global paleoceanography. Journal of Geophysical Research, 82, 3843–3860. Kitoh, A. 2004. Effects of mountain uplift on East Asian summer climate investigated by a coupled atmosphereocean GCM. Journal of Climatology, 17, 783–802. Molnar, P., England, P. & Martinod, J. 1993. Mantle dynamics, uplift of the Tibetan Plateau, and the Indian Monsoon. Reviews of Geophysics, 31, 357– 396. Najman, Y. 2006. The detrital record of orogenesis: a review of approaches and techniques used in the Himalayan sedimentary basins. Earth-Science Reviews, 74, 1– 72. Prell, W. L. & Kutzbach, J. E. 1992. Sensitivity of the Indian Monsoon to forcing parameters and implications for its evolution. Nature, 360, 647– 652. Ramstein, G., Fluteau, F., Besse, J. & Joussaume, S. 1997. Effect of orogeny, plate motion and land-sea distribution on Eurasian climate change over the past 30 million years. Nature, 386, 788– 795. Raymo, M. E. 1994. The Himalayas, organic-carbon burial, and climate in the Miocene. Paleoceanography, 9, 399–404. Raymo, M. E. & Ruddiman, W. F. 1992. Tectonic forcing of Late Cenozoic climate. Nature, 359, 117–122. Royden, L. H., Burchfiel, B. C. & Van Der Hilst, R. D. 2008. The geological evolution of the Tibetan plateau. Science, 321, 1054–1058. Tada, R. 2004. Onset and evolution of millennial-scale variability in the Asian monsoon and its impact on paleoceanography of the Japan Sea. In: Clift, P. D., Kuhnt, W., Wang, P. & Hayes, D. (eds) Continent– Ocean Interactions within East Asian Marginal Seas. American Geophysical Union, Washington D.C., Geophysical Monograph, 149, 283–297. Tapponnier, P., Xu, Z. Q., Roger, F., Meyer, B., Arnaud, N., Wittlinger, G. & Yang, J. S. 2001. Geology – oblique stepwise rise and growth of the Tibet plateau. Science, 294, 1671–1677. Thiede, R. C., Bookhagen, B., Arrowsmith, J. R., Sobel, E. R. & Strecker, M. R. 2004. Climatic control on rapid exhumation along the Southern Himalayan Front. Earth and Planetary Science Letters, 222, 791– 806. von der Heydt, A. & Dijkstra, H. A. 2006. Effect of ocean gateways on the global ocean circulation in the late Oligocene and early Miocene. Paleoceanography, 21, doi: 10.1029/2005PA001149. Webster, P. J., Magana, V. O., Palmer, T. N., Shukla, J. R. A., Tomas, M., Yanai, Y. & Yasunari, T. 1998. Monsoons: processes, predictability, and the prospects for prediction, in the TOGA decade. Journal of Geophysical Research, 103, 14 451– 14 510. Willett, S. D. 1999. Orogeny and orography: the effects of erosion on the structure of mountain belts. Journal of Geophysical Research, 104, 28 957– 28 981. Zheng, H., Powell, C. M., Rea, D. K., Wang, J. & Wang, P. 2004. Late Miocene and mid-Pliocene enhancement of the east Asian monsoon as viewed from the land and sea. Global and Planetary Change, 41, 147– 155.

Marine records of East Asian monsoon variability over the past 5 Ma MIN-TE CHEN*, YUAN-PIN CHANG, PAI-SEN YU & LIANG-JIAN SHIAU Institute of Applied Geosciences, National Taiwan Ocean University, Keelung 20224, Taiwan *Corresponding author (e-mail: [email protected]) Abstract: Marine sedimentary cores retrieved from the western Pacific provide important clues for deciphering how the East Asian Monsoon (EAM) system has evolved during the past 5 Ma. Here we briefly review some recent progress on the reconstructions of the EAM based on marine SST (sea surface temperature), SSS (sea surface salinity), and productivity records from the SCS (South China Sea) and ECS (East China Sea) and their implications for EAM evolution and variability on tectonic, orbital and millennial timescales. This review highlights the importance of high resolution sampling on giant marine cores (such as cores collected with the International Marine Past Global Change, IMAGES program) that provide opportunities for better defining the timing and amplitude of the EAM variability expressed in marine records. We also discuss possible future directions of EAM palaeoclimatic and palaeoceanographic studies that require development of multiple new marine EAM proxies and a comparison of the marine records with the stalagmite records on land.

The East Asian Monsoon (EAM) is an important climate modulator with links to the climate systems of the high latitudes, as well as major elements of the general atmospheric circulation system dominating the tropics, such as ENSO (El Nin˜o Southern Oscillation) (Wang et al. 2003; Wang & Ding 2008). Climatologically, monsoons are driven by differential land –sea sensible heating in different seasons. The monsoon-dominated regions are the most convectively active areas on Earth and account for the majority of global atmospheric heat and moisture transport (Webster 1987). The EAM, as a modern active component of the climate system of East Asia, determines most of the variability in the SST (sea surface temperature) and SSS (sea surface salinity) of the western tropical Pacific over seasonal (Fig. 1) to interannual timescales (Wang et al. 2003). The EAM thus is characterized by two seasonal components: namely East Asian Winter Monsoon (EAWM) and East Asian Summer Monsoon (EASM). Understanding how the EAM might have evolved over longer timescales (.Ma) is highly relevant to the future development of human societies located in the tropics and subtropics, and is still a scientifically challenging question. Marine sedimentary records provide much information on EAM variability over longer timescales (tectonic, orbital, and millennial) (Prell et al. 1992; Clemens et al. 2008; Clift et al. 2008). Most palaeoceanographic and palaeoclimatological data and records about the long-term evolution and variability of the EAM have come from studies on

marine sedimentary cores from stable and high sedimentation rate areas in western Pacific marginal seas [e.g. the South China Sea (SCS) and East China Sea (ECS)] (Chen & Huang 1998; Chang et al. 2009). Recent IMAGES (International Marine Past Global Change Study) and ODP (Ocean Drilling Program) cruises to the western Pacific represent such efforts for collecting the best marine geological archives to reconstruct longterm EAM evolution and variability (Fig. 2). Although SST and SSS have been used extensively as EAM proxies in these marine core studies, the interpretation of SST and SSS appear to be complicated by many other climatic factors. The fluctuations in the Oyashio/Kuroshio patterns (Sawada & Handa 1998), stationary rain fronts (Mei-Yu), and the intertropical convergence zone (ITCZ) may all be responsible for determining the annual to interannual, and therefore long-term variability, of the EAM (Wang et al. 2003). Moreover, some farreaching factors such as the Walker Circulation in the Pacific, cross-equatorial flow driven by a geographic forcing of Australia, and eustatic sea-level change also have been suggested to explain the long-term EAM variability (An 2000). Distinguishing this complicated expression of the EAM in marine records relies on the careful evaluation and usage of proxies, establishment of more latitudinal and longitudinal transects of records, and comparison of the marine records with the terrestrial records. In this paper, we review the current marine evidence of EAM evolution and variability at

From: Clift, P. D., Tada, R. & Zheng, H. (eds) Monsoon Evolution and Tectonics –Climate Linkage in Asia. Geological Society, London, Special Publications, 342, 5–15. DOI: 10.1144/SP342.2 0305-8719/10/$15.00 # The Geological Society of London 2010.

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Fig. 1. EAM-related January and July SST (Sea Surface Temperature) and SSS (Sea Surface Salinity) in the western Pacific (data derived from MODAS regressed satellite synthetic T and S profile) (data source: Naval Research Lab Global NCOM; http://www7320.nrlssc.navy.mil/modas/).

tectonic, orbital and millennial timescales. Most of the evidence presented here is based primarily on recent investigations on IMAGES and ODP cores retrieved from the SCS and ECS. As the EAM is a climatic feature that mainly affects surface ocean hydrography and productivity, what

we review here are mostly high quality SST, SSS, and surface productivity records for the EAM. We conclude by proposing future directions for exploring more new proxies and new records to improve our understanding of the evolution and variability of the EAM.

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Fig. 2. Locations of EAM marine core and cave sites discussed in the paper.

Tectonic evolution The uplift of the Himalayan-Tibetan Plateau has long been thought to be a primary driver for the long-term evolution of the EAM on tectonic timescales (An et al. 2001). Evidence for the long-term evolution of the EAM comes primarily from studies using loess, lake and marine sediment records, whereas this review focuses on synthesizing the

EAM history of the past 5 Ma based on data from published and on-going works of marine sediment cores, mostly from the SCS and ECS. In loess studies for the EAM evolution, the EAM is thought to have been initiated during the much earlier time of c. 22 Ma ago (Guo et al. 2002). The proposed timing for the early initiation of the EAM is based on observations on old aeolian layers that are characterized by alternation of brownish loess and

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reddish soils, and is 14 Ma older than the timing suggested by previous studies on the Loess Plateau (Sun et al. 1998a, b). Marine records of the EAM are still too short to document the long-term initiation of the EAM. The microfossil evidence suggests that the increased strength of summer monsoon upwelling or thermocline deepening might have evolved in the SCS since c. 8–6 Ma (M. Chen et al. 2003; Li et al. 2004). Although the microfossil data appear to be subject to larger interpretation uncertainty for EAM-induced hydrographic conditions, the suggested timing echoed what has been proposed for the increased strength of the EASM observed from loess sequences (Sun et al. 1998a, b). The development of the initial EASM is defined as the first stage of EASM evolution, which is thought to respond to the enhanced aridity in the Asian interior environment since that time (An et al. 2001). The further enhancement of the EASM is identified in marine and terrestrial records of between 3.6– 2.6 Ma (Tian et al. 2004; Fig. 3). Planktic and benthic isotope records from ODP Site 1143 suggest a decrease of SSS in the SCS since 3.5 Ma (Tian et al. 2004). The possible freshening and increased EASM precipitation since 3.5 Ma are corroborated by evidence for increased strength and alternation of humidity and aridity in the eastern Loess Plateau (An et al. 2001) and the western

Kunlun Mountains (Zheng et al. 2000), and for increased variability of aeolian fluxes in the North Pacific (Rea et al. 1998). Astronomically-tuned marine oxygen isotope records (Raymo et al. 2006) suggest that the North Hemisphere glaciation was initiated around 2.75 Ma. The cooling of North Hemisphere continents by the glaciation increased the temperature contrast between the East Asia and the western Pacific, therefore favoured the development of stronger EASM. However, the timing of the second stage of the strengthened EASM appears to be too early for the initiation of the North Hemisphere glaciation at c. 2.75 Ma. Thus the uplift of the Himalayan-Tibetan Plateau might be involved in determining the timing of the stronger EASM intensity and variability during this time interval. Moreover, a recent synthesis of EASM marine records suggests that other climatic factors should be also important in determining the timing of the EASM evolution during the time interval 3.6–2.6 Ma (Clemens et al. 2008). Marine and tectonic evidence has recently become available that indicates that the gradual change in the long-term state of atmospheric and oceanic circulations in the tropical Pacific might be responsible for the evolution of the EASM since 5 Ma (Haug & Tiedemann 1998; Cane & Molnar 2001; Ravelo et al. 2004; Wara et al. 2005; Raymo et al. 2006) (Fig. 4). In the modern

4.6 Ma Panama Isthmus Closure 4–3 Ma Indonesian Seaway Closure Pre-conditioning of Pacific ‘Walker Circulation’ – increase of latitudinal transport of moisture/heat and tropical A-O interaction

3.6–2.6 Ma Mountain Uplifting Further intensification of summer EAM by land sensible heating – tectonic evolution of EAM

2.75 Ma Northern Hemisphere Glaciation Further intensification of winter EAM by increase of high-low latitude thermal gradient – orbital variability of EAM

1.7 Ma Final Establishment of Pacific ‘Walker Circulation’ Further intensification of moisture/heat transport – orbital and more millennial variability of EAM – possibly amplified by tropical A-O interaction Fig. 3. The sequence of tectonic and climatic events that could be responsible for intensifying the EAM in the past 5 Ma.

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Fig. 4. (a) Carbonate (wt%); (b) TOC (wt%); (c) opal (wt%); (d) C37 alkenones (mg/g); (e) CPI (a composite productivity index); (f) alkenone SST (oC) (in red) and d18O (in blue) records of MD972142 from the southwestern SCS. The dashed line indicates a long-term trend of increased productivity since the middle Pleistocene. The alkenone SST a d18O variations show nearly in-phase relationships. (Reproduced from Shiau et al. 2008.)

pattern of atmospheric circulation in the tropical Pacific, the EASM continues as a division of the trade winds blowing from the eastern equatorial Pacific (Webster 1987). The intensification of the EASM must partly reflect the basin-scale climate conditions in the Pacific. Changes toward modern climate conditions analogous to ‘La Nin˜a’ in the Pacific favour initiation of a stronger EASM (Meehl 1987; Yasunari 1991), as the ‘La Nin˜a’ condition maintains or even expands the extent of western Pacific warm pool, that in turn increases the transport of heat and moisture from the tropical western Pacific to the interior of the Asia, analogous to the climate conditions produced by stronger

EASM. During 4.6– 3 Ma, the gradual closure of the Panama Isthmus and Indonesian Seaways might have provided a ‘pre-condition’ for more tilted thermocline or SST gradients across the equatorial Pacific, and in turn shifted the Pacific climate into more ‘La Nin˜a’-like conditions that were involved in the early development of the EASM (Haug & Tiedmann 1998; Cane & Molnar 2001). Consequently, the strengthened EASM induced the increase of surface water freshness in the SCS (Tian et al. 2004) and more frequent alternations of the humid and dry intervals observed in the loess profiles in central China of c. 3.6 –2.6 Ma (An et al. 2001). The 1 –1.5 Ma delay of EASM

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intensification with respect to the closure of the Panama Isthmus and Indonesian Seaways might reflect the time required for the EASM to respond fully to the change in Pacific circulation patterns, or due to the age uncertainty of the tectonic events. The final establishment of the ‘La Nin˜a’conditions or western Pacific warm pool has been suggested to be during c. 1.7 Ma; the evidence comes from a gradient reconstruction of foraminiferal isotope records from ODP Site 847, 851 (eastern equatorial Pacific) and Site 806 (western equatorial Pacific) (Cannariato & Ravelo 1997; Ravelo et al. 2004; Wara et al. 2005). Since c. 1.7 –1 Ma, a greater supply of moisture and heat from the tropical Indian and Pacific to the interior of the Asia, along with the initiation of Northern Hemisphere glaciation, have both played roles in maintaining the stability of more frequent and intensified EASM and EAWM, respectively (Fig. 4). The timing of intensification for the EASM since c. 4.6–3 Ma suggested by the evidence presented above is consistent with what is observed from the increased wetness on the mountain range of the Himalayas, based on new chemical weathering records reconstructed from ODP Site 1148 (SCS) and Site 718 (Bengal Fan) (Clift et al. 2008), and therefore indicates a complex cause-and-effect relationship between monsoons and land erosion (deformation).

Orbital and millennial scale variability Since c. 1.7–1 Ma, the most dominant factors determining the timing and strength of the EAM variability over orbital to millennial timescales appear to be related to insolation and glaciation (Huang et al. 1997a, b; Chen & Huang 1998; Chen et al. 1999). The EAM variability over orbital to millennial timescales has been studied extensively through the use of IMAGES and ODP records from the SCS (see the review in Wang et al. 2005). The EAWM is particularly well expressed in the SCS records, with pronounced cooling, deep mixing, and possibly high productivity during glacial times when the EAWM is stronger (Huang et al. 1997a, b). Precession and millennial-scale variations of the EAWM have also been reported from records retrieved from high sedimentation rate sites in the SCS (Chen & Huang 1998; Chen et al. 1999). A long record from ODP Site 1143 which covers the past 0.8 Ma (Tian et al. 2005) clearly indicates a contrast of glacial and interglacial variations in surface ocean mixing, productivity, and continental aridity in the SCS. All evidence accumulated from ODP Leg 184 SCS records thus suggested that the EAM would have been an important agent in controlling this ‘two-mode’ EAM hydrographic change in the

SCS during the past 1 Ma. Stronger EAWM induces cooling and high productivity in the ocean, and aridity on land; weaker EAWM occurs with warming and low productivity in the ocean, and humidity on land. With the success of more recent IMAGES programs in the SCS (Chen et al. 1998; Bassinot et al. 2002), high- to ultra-high resolution records have become available for studying EAM variability on orbital to millennial-scales (and possibly extended to centennial scales). The high- to ultrahigh resolution EAM records provide better opportunities to define the timing and amplitude of the EAM variability. The large volume of IMAGES core sediments also permits multiple proxy analyses of palaeoceanography through the use of samples of the same depth. One example of such multiple proxy studies is a study of IMAGES core MD972142 located at the southwestern SCS (Shiau et al. 2008) (Fig. 4). In MD972142, SSTs have been reconstructed by the use of alkenone unsaturation indicators, which could be compared to SSTs estimated from planktic foraminifer fauna assemblage data from the same core (M.-T. Chen et al. 2003) that reveal interesting differences in the timing and amplitude of SST reconstructions by different methods. While the maximum of alkenone SST of MD972142 shows nearly in-phase relationship with ice volume minimum over precession frequency band (Shiau et al. 2008), the maximum fauna SST lags significantly relative to ice volume minimum over the frequency band (M.-T. Chen et al. 2003), but within the same phase range as maximum summer monsoon reported based on Indian Ocean records (Clemens et al. 2008). The differences in the timing of the alkenone and fauna SST reconstructions may be attributed to the seasonality, preservation, and sensitivity of the SST proxies that are used. Productivity variations in MD972142 have been estimated through the use of a composite of multiple proxies – concentrations of carbonate, opal, total organic carbon, and alkenones (Fig. 4) – for reconciling the discrepancies in timing and amplitude shown in these different proxies. The composite productivity index, which is the first eigenvector which explains 37% of the total variance of these multiple proxies in MD972142 shows increased productivity during glacial intervals of the past c. 1 Ma (Shiau et al. 2008) (Fig. 4). The high glacial productivity in the SCS appears to be in contradiction to the high productivity during interglacials, as suggested by biogenic-related element records from the southern SCS (Wei et al. 2003). The discrepancy in the productivity interpretation might be minimized by careful evaluation of terrestrial sediment fluxes into the SCS and their influences on the biogenic components preserved in SCS sediments, as high

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terrestrial sediment fluxes favour good preservation of the productivity proxies, and then may bias the pattern of productivity reconstructions. Quantification of terrestrial sediment fluxes against productivity reconstructions in the same core may help minimize the preservation bias. More interestingly, the productivity variations show long-term increasing trends since the mid-Brunhes, most noticeably from c. 330 ka. Uplifting of the Tibetan Plateau has accelerated since that time (An 2000) and may be responsible for the strengthening of the EAWM and long-term increase in productivity in the SCS. Interpreting the hydrographic changes caused by EAM in SCS marine records of the past 1 Ma is complicated by the eustatic sea level changes which control the proximity of river mouths, more or less semi-enclosed basins, and continental weathering and erosion that may bias the accuracy of hydrographic proxies for the EAM. To minimize the complicated effects induced by the eustatic sea level changes, a ‘gradient approach’ that involves the use of multiple or a transect of records is encouraging in identifying hydrographic changes related to EAWM variability. An example of studies that have adopted the strategy is found in a reconstruction of fauna SST gradients for the eastern and western SCS (IMAGES MD972142 and MD012394) (Yu et al. 2006) (Fig. 5). This study assumes that the winter SST gradients between the eastern and western SCS could be used to monitor the variations of the past EAWM intensity. The assumption is based on modern observation of a large SST gradient occurring between eastern and western SCS during maximum winter seasons (Fig. 1). Following this assumption, Yu et al. (2006) have selected two records, which are located in a sharp contrast of hydrographic conditions, in which the EAWM drives cold and strong mixing conditions in the western SCS and maintains a warm and deep-thermocline in the southeastern SCS. Changes in the planktic foraminifer fauna assemblages of these two cores were used to evaluate the dynamics of surface hydrographic conditions related to the EAWM. SST and fauna hydrographic gradients reconstructed in this study show a dominance of 23 ka precession cycles over the past 135 ka. Large SST and hydrographic gradients imply a stronger EAWM (or weaker EASM) at 22, 45, 71, 94, and 116 ka. In the precession cycles, the SST and hydrographic gradients increased at the December 21 perihelion, which is associated with seasons of minimum Northern Hemisphere winter and maximum South Hemisphere summer insolation. Cross-spectral analyses between the SST and hydrographic gradients with the precession indicate coherent and in-phase relationships. Considering the seasonal cycles during the December 21

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perihelion are in accord with the maximum summer insolation in the Southern Hemisphere, this study reports nearly in-phase relationship between maximum EAWM in the SCS and the orbital configuration for maximizing sensible heating over Southern Hemisphere continents during the summer – a condition which favours the development of Australian summer monsoon. Based on the observation, this study suggests that the EAM variability in the SCS was dominated by winter monsoon dynamics for the last 135 ka in response to an inter-hemispheric forcing that drives a cross-equatorial monsoon flow from the interior of Asia to Australia. The EAM variability is not only expressed in the SCS. Recent studies that take advantage of high sedimentation rate cores from the ECS (mainly from the Okinawa Trough) also indicate that the significant orbital to millennial-scale variations in surface marine productivity are closely related to the precipitation of the EASM on land (Chang et al. 2009; Fig. 6). The variability shown in the records (IMAGES MD012404) on orbital timescales indicates that high TOC intervals coincide with the increases of boreal late spring–early autumn insolation driven by precession cycles (c. 21 ka), implying a strong connection to the variations in monsoons. Moreover, this study observes varied timing of monsoon productivity responses, in which some monsoon maxima corresponding to May or September insolation maxima. The varied timing could be explained by the fact that the other climate mechanisms dominated in the western Pacific such as Mei-Yu or typhoon are also important in determining the timing and amplitude of EASM, as Mei-Yu and typhoon bring heavy precipitation in May and September, respectively in modern western Pacific climate. This study also observes possibly nearly synchronous, millennialscale changes in the ECS surface hydrography (mainly driven by salinity changes but also by temperature effects) and productivity coincident with monsoon events in the Hulu/Dongge stalagmite isotope records (Wang et al. 2008). Relatively high productivity and fresh surface water correspond to interstadials and low productivity and saline surface water occur during stadials. The high productivity and increased freshening correlate with high monsoon intensity in interstadials, suggesting that the millennial-scale changes in monsoon hydrography and productivity in the ECS have been remarkable and persistent features over the past 100 000 years.

Conclusions and future directions We have reviewed the current marine evidence for documenting and understanding EAM

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Fig. 5. Faunal factor, hydrographic, and SST gradients shown in eastern (MD972142) and western (MD012394) SCS cores. The figure shows the differences between the two records: (a) warm water species factor; (b) cold water species factor; (c) planktic foraminifer hydrographic index (PFHI) as defined by the ratio of warm/cold water species factors (Yu et al. 2006); (d) SST estimated by MAT (Modern Analogue Technique); and (e) planktic foraminifera d18O (MD972142 in red and MD012394 in blue), and plotted against age and compared to precession index variations (in green). A high precession index (Pmax) indicates an orbital configuration of the December 21 perihelion and a low (Pmin) index indicates the June 21 perihelion. Large faunal factor, hydrographic, and SST gradients are associated with a high precession index. (Reproduced from Yu et al. 2006.)

variability on tectonic, orbital, and millennial timescales. Most of the high quality marine evidence for studying the EAM history comes from cores retrieved from the SCS and ECS. Marine records

for EAM evolution and variability available at present are too short to cover the geological range that terrestrial records cover. Increased drilling (or coring) for long records at high sedimentation rate

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Fig. 6. Productivity proxies (TOC and opal) measured from MD012404 (ECS) for evaluating monsoon precipitation and productivity changes in the Okinawa Trough over the past 100 000 years. These data are compared to a 608N Northern Hemisphere July insolation (in green) (Berger 1978), and the d18O record from Hulu-Sanbao stalagmites from East China (Wang et al. 2008). Age controls of the records are presented in Chang et al. (2009). The dashed lines are correlations of millennial-scale warm events. (Reproduced from Chang et al. 2009.)

sites in the western Pacific marginal seas is necessary to obtain better temporal and spatial coverage for the EAM evolution and variability from c. 22 Ma. Candidates for such drilling or coring that have been less explored include the southern part of the SCS and the ECS. All previous and recent studies that address EAM evolution and variability have suggested that the glacial stages are characterized by stronger EAWM, while the interglacial stages are more affected by the EASM. Precession cycles have been the most dominant variations in most late Quaternary marine EAM records. The timing of the EAM variations based on SST reconstructions shows interesting proxy-dependent discrepancies (Steinke et al. 2008). Resolving the discrepancies must rely on more systematic efforts in evaluating

multiple proxies of SSTs (fauna, alkenone and Mg/Ca) by compiling core top (Chen et al. 2005) and sediment trap data sets, or by testing isotopic or geochemical proxies on individual microfossil species (such as Globigerinoides ruber (white) morphotypes, Lo¨wemark et al. 2005; Steinke et al. 2005). The timing of maximum surface marine productivity in the SCS also appears to be contentious in reconstructions using biogenic sediment components (i.e. carbonate, opal, and alkenone contents) and element ratios (i.e. Ba/Ca), whereas the productivity maxima in the ECS coincide more closely to the precession minima (or interglacial maxima). Millennial-scale hydrographic and productivity variations appear to be a dominant feature in most high resolution EAM records from the SCS and ECS, and the timing of the millennial variations,

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in most cases, shows warm, minimal surface mixing (or less salinity) which indicates a stronger EASM in interstadials, and cold and more surface mixing (stronger EAWM) in stadials. The timing of these variations is shown to be in-phase with what has been established precisely in stalagmite EASM records (Wang et al. 2008). Future comparisons using stalagmite and marine records may provide exciting opportunities for examining more regional and more precisely dated patterns in past EAM evolution and variability. We thank the National Science Council and the National Taiwan Ocean University, Taiwan, for continuous support of Taiwan’s IMAGES palaeoceanographic programs (1996–2008) and for the core repository and laboratory facilities of the National Center for Ocean Research. We thank the organization for the IGCP-476 workshops ‘Monsoon evolution and tectonics: climate linkage in Asia’ that made the proposal of this special issue publication possible.

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Chen, M.-T., Wang, C.-H., Huang, C.-Y., Wang, L. & Sarnthein, M. 1999. A late Quaternary planktonic foraminifer faunal record of rapid climatic changes from the South China Sea. Marine Geology, 156, 85–108. Chen, M.-T., Shiau, L.-J., Yu, P.-S., Chiu, T.-C., Chen, Y.-G. & Wei, K.-Y. 2003. 500 000-year records of carbonate, organic carbon, and foraminiferal seasurface temperature from the southeastern South China Sea (near Palawan Island). Palaeogeography, Palaeoclimatology, Palaeoecology, 197, 113– 131. Chen, M.-T., Huang, C.-C., Pflaumann, U., Waelbroeck, C. & Kucera, M. 2005. Estimating glacial western Pacific sea-surface temperature: methodological overview and data compilation of surface sediment planktic foraminifer faunas. Quaternary Science Reviews, 24, 1049– 1062. Chen, M., Wang, R., Yang, L., Han, J. & Lu, J. 2003. Development of east Asian summer monsoon environments in the late Miocene: radiolarian evidence from Site 1143 of ODP Leg 184. Marine Geology, 201, 169–177. Clemens, S. C., Prell, W. L., Sun, Y., Liu, Z. & Chen, G. 2008. Southern Hemisphere forcing of Pliocene d18O and the evolution of Indo-Asian monsoons. Paleoceanography, 23, PA4210, doi: 10.1029/ 2008pa001638. Clift, P. D., Hodges, K. V., Heslop, D., Hannigan, R., Van Long, H. & Calves, G. 2008. Correlation of Himalayan exhumation rates and Asian monsoon intensity. Nature Geosciences, 1, 1 –6. Guo, Z. T., Ruddiman, W. F. et al. 2002. Onset of Asian desertification by 22 Myr ago inferred from loess deposits in China. Nature, 416, 159–163. Haug, G. H. & Tiedemann, R. 1998. Effect of the formation of the Isthmus of Panama on Atlantic Ocean thermohaline circulation. Nature, 393, 673–676. Huang, C.-Y., Liew, P.-M. et al. 1997a. Deep sea and lake records of the Southern Asian paleomonsoons for the last 25 thousand years. Earth and Planetary Science Letters, 146, 59–72. Huang, C.-Y., Wu, S.-F., Zhao, M., Chen, M.-T., Wang, C.-H., Tu, X. & Yuan, P. B. 1997b. Surface ocean and monsoon climate variability in the South China Sea since the last glaciation. Marine Micropaleontology, 32, 71–94. Li, B., Wang, J. et al. 2004. South China Sea surface water evolution over the last the 12 Myr: a southnorth comparison from Ocean Drilling Program Sites 1143 and 1146. Paleoceanography, 19, PA1009, doi: 1010.1029/2003PA000906. Lo¨wemark, L., Hong, W.-L., Yui, T.-F. & Hung, G.-W. 2005. A test of different factors influencing the isotopic signal of planktonic foraminifers in surface sediments from the northern South China Sea. Marine Micropaleontology, 55, 49–62. Meehl, G. A. 1987. The annual cycle and interannual variability in the tropical Pacific and Indian Ocean regions. Monthly Weather Review, 115, 27– 50. Prell, W. L., Murray, D. W., Clemens, S. C. & Anderson, D. M. 1992. Evolution and variability of the Indian Ocean summer monsoon: evidence from the western Arabian Sea drilling program.

EAST ASIAN MONSOON VARIABILITY In: Duncan, R. A., Rea, D. K., Kidd, R. B., von Rad, U. & Weissel, J. K. (eds) Synthesis of Results from Scientific Drilling in the Indian Ocean, Geophysical Monograph 70. America Geophysical Union, Washington, DC, 447– 469. Ravelo, A. C., Andreasen, D. H., Lyle, M., Olivarez Lyle, A. & Wara, M. W. 2004. Regional climate shifts caused by gradual global cooling in the Pliocene epoch. Nature, 429, 263–267. Raymo, M. E., Lisiecki, L. E. & Nisancioglu, K. H. 2006. Plio-Pleistocene ice volume, Antarctic climate, and the global d18O record. Science, 313, 492– 495. Rea, D. K., Snoeckx, H. & Joseph, L. H. 1998. Late Cenozoic eolian deposition in the North Pacific: Asian drying, Tibetan uplift, and cooling of the northern hemisphere. Paleoceanography, 13, 215–224. Sawada, K. & Handa, N. 1998. Variability of the path of the Kuroshio ocean current over the past 25 000 years. Nature, 392, 592– 594. Shiau, L.-J., Yu, P.-S. et al. 2008. Sea surface temperature, productivity, and terrestrial flux variations of the southeastern South China Sea over the past 800 000 years (IMAGES MD972142). Terrestrial, Atmospheric and Oceanic Sciences, 19, 363– 376. Steinke, S., Chiu, H.-Y., Yu, P.-S., Shen, C.-C., Lo¨wemark, L., Mii, H.-S. & Chen, M.-T. 2005. Mg/Ca ratios of two Globigerinoides ruber (white) morphotypes: implications for reconstructing past tropical/subtropical surface water conditions. Geochemistry Geophysics Geosystems, 6, Q11005, doi: 11010.11029/12005GC000926. Steinke, S., Kienast, M., Groeneveld, J., Lin, L.-C., Chen, M.-T. & Rendle-Bu¨hring, R. 2008. Proxy dependence of the temporal pattern of deglacial warming in the tropical South China Sea: toward resolving seasonality. Quaternary Science Reviews, 27, 688–700. Sun, D., Shaw, J., An, Z., Cheng, M. & Yue, L. 1998a. Magnetostratigraphy and paleoclimatic interpretation of a continuous 7.2 Ma Late Cenozoic eolian sediments from the Chinese Loess Plateau. Geophysical Research Letters, 25, 85–88. Sun, D., An, Z., Shaw, J., Bloemendal, J. & Sun, Y. 1998b. Magnetostratigraphy and palaeoclimatic significance of Late Tertiary aeolian sequences in the Chinese Loess Plateau. Geophysical Journal International, 134, 207– 212.

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Tian, J., Wang, P. & Cheng, X. 2004. Development of the East Asian monsoon and Northern Hemisphere glaciation: oxygen isotope records from the South China Sea. Quaternary Science Reviews, 23, 2007– 2016. Tian, J., Wang, P., Cheng, X., Wang, R. & Sun, X. 2005. Forcing mechanism of the Pleistocene east Asian monsoon variations in a phase perspective. Science in China, 48, 1708–1717. Wang, B. & Ding, Q. 2008. Global monsoon: dominant mode of annual variation in the tropics. Dynamics of Atmospheres and Oceans, 44, 165 –183, doi: 10.1016/j.dynatmoce.2007.05.002. Wang, B., Clemens, S. C. & Liu, P. 2003. Contrasting the Indian and East Asian monsoons: implications on geologic timescales. Marine Geology, 201, 5 –21, doi: 10.1016/S0025-3227(03)00196-8. Wang, P., Clemens, S. C. et al. 2005. Evolution and variability of the Asian monsoon system: state of the art and outstanding issues. Quaternary Science Reviews, 24, 595–629, doi: 10.1016/j.quascirev.2004. 10.002. Wang, Y., Cheng, H. et al. 2008. Millennial- and orbital-scale changes in the East Asian monsoon over the past 224,000 years. Nature, 451, 1090–1093. Wara, M. W., Ravelo, A. C. & Delaney, M. L. 2005. Permanent El Nin˜o-like conditions during the Pliocene warm period. Science, 309, 758– 761. Webster, P. J. 1987. The elementary monsoon. In: Fein, J. S. & Stephens, P. L. (eds) Monsoons. John Wiley & Sons, New York, 3– 32. Wei, K.-Y., Chiu, T.-C. & Chen, Y.-G. 2003. Toward establishing a maritime proxy record of the East Asian summer monsoons for the late Quaternary. Marine Geology, 201, 67–79. Yasunari, T. 1991. The Monsoon Year – A new concept of the climatic year in the tropics. Bulletin of the American Meteorological Society, 72, 1331–1338. Yu, P.-S., Huang, C.-C., Chin, Y., Mii, H.-S. & Chen, M.-T. 2006. Late Quaternary East Asian Monsoon variability in the South China Sea: evidence from planktonic foraminifera faunal and hydrographic gradient records. Palaeogeography, Palaeoclimatology, Palaeoecology, 236, 74–90. Zheng, H., Powell, C. M., An, Z., Zhou, J. & Dong, G. 2000. Pliocene uplift of the northern Tibetan Plateau. Geology, 28, 715–718.

Orbital and suborbital variability in the equatorial Indian Ocean as recorded in sediments of the Maldives Ridge (ODP Hole 716A) during the past 444 ka SOMA DE1, SUDIPTA SARKAR2 & ANIL K. GUPTA1* 1

Department of Geology and Geophysics, Indian Institute of Technology, Kharagpur 721 302, India

2

Department of Geology, Rajiv Gandhi Institute of Petroleum Technology, Rae Bareli 229316, U.P., India *Corresponding author (e-mail: [email protected])

Abstract: This study is aimed at understanding past 444 ka record of climate variability in the equatorial Indian Ocean using high resolution records of planktic and benthic foraminifera and pteropods from Ocean Drilling Program Hole 716A, Maldives Ridge, southeastern Arabian Sea. In total, 892 samples of 10 cm3 volume from 444 ka old sequence were analysed at 1 cm intervals to generate census data of the foraminiferal fauna and pteropods. The percent and detrended time series of mixed-layer species Globigerinoides ruber and Globigerinoides sacculifer and thermocline species Neogloboquadrina dutertrei, benthic foraminifera Cymbaloporetta squammosa, Sphaeroidina bulloides and Uvigerina proboscidea, and pteropods from ODP Hole 716A reveal significant changes in wind intensity during the past 444 ka. An abrupt decrease in the Cymbaloporetta squammosa population at c. 300 ka (across MIS 8/9) suggests a weakening of equatorial wind intensity, which could be linked to Indian monsoon and may have driven pronounced changes in the oxygen minimum zone in the Maldivian region. These changes were contemporaneous with the Mid-Brunhes Climatic Event, the beginning of aridity in the Indonesian-Australian region and the onset of a humid phase in equatorial East Africa as observed in several oceanic and continental records. This strengthens a connection between equatorial Indian Ocean wind intensity, the Indian monsoon and Indonesian-Australian-African climates. Supplementary material: Percentages of benthic and planktic foraminifera and pteropods used in the present study are available at http://www.geolsoc.org.uk/SUP18413

The northwestern Indian Ocean climate is marked by regular rhythms of seasonal changes in wind character and rains over the South Asian landmass, known as summer or southwest (SW) and winter or northeast (NE) monsoons. Thermal contrast between land and sea owing to sensible heating of the Asian landmass and latent heat released from precipitation due to cross-equatorial moisture transport, drive the monsoonal circulation (Webster 1987). During summer (June to September), due to heating of the Asian landmass the winds are southwesterly, intense and moist, carrying moisture to the South Asian landmass. In winter (November to February), the land is colder than sea and the wind direction reverses, blowing from NE to SW. For the inhabitants of South Asia, the rainy season means summer monsoon season because most of the rainfall occurs during this time. It is crucial for agricultural production, water needs and the economic development of the South Asian societies. The regularity and importance of monsoonal rains has led to the evolution of traditional cultures and the proverbs of the people of South Asia, in general,

and of India, in particular. The SW monsoon variability drives significant changes in the biogeochemical character of the northern Indian Ocean, with stronger summer monsoon winds causing widespread upwelling, high surface productivity and proliferation of distinct fauna and flora in different parts of the northwestern Indian Ocean. The NE monsoon winds are dry, weak and variable, causing weak upwelling cells and patchy surface productivity in the region. The high surface productivity during SW monsoon season enhances biogenic sediment accumulation in the Indian Ocean, providing us with the best proxy record of summer monsoon variability at centennial and millennial scales. Extreme events in the monsoon cause severe floods and crop failures impacting human life and thus economy of the South Asian region that houses about 60% of the world’s population (Webster et al. 1998). The recent increase in global surface temperature has been implicated in bringing significant change in the monsoon behaviour with increased frequency and magnitude of extreme rain events in India over the past 50 years

From: Clift, P. D., Tada, R. & Zheng, H. (eds) Monsoon Evolution and Tectonics –Climate Linkage in Asia. Geological Society, London, Special Publications, 342, 17– 27. DOI: 10.1144/SP342.3 0305-8719/10/$15.00 # The Geological Society of London 2010.

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(Goswami et al. 2006). Thus, understanding small scale variations in the Indian summer monsoon has both scientific and societal impact. Recent studies have suggested a link between Atlantic Multidecadal Oscillation (AMO) (Goswami et al. 2006; Lu et al. 2006) or North Atlantic thermohaline circulation (Zhang & Delworth 2005) and interannual variability in the Indian summer monsoon. These studies suggest a potential role of the Atlantic Ocean in the low frequency fluctuation of the Indian monsoon. The positive AMO phase with anomalous warm North Atlantic and cold South Atlantic has been suggested as a possible mechanism causing stronger Indian and East Asian monsoons (Lu et al. 2006). It has also been suggested that the Indian monsoon responds to changes in El Nin˜o-Southern Oscillation (ENSO) (Overpeck et al. 1996). However, historical records indicate a weakened relationship between the Indian monsoon and ENSO owing to a shift in the Walker circulation and/or an increase in surface temperature over Eurasia (Krishna Kumar et al. 1999). A recent model study suggests that El Nin˜o events with the warmest sea surface temperature anomalies in the central equatorial Pacific are more effective in causing monsoon failures (Krishna Kumar et al. 2006). The role of equatorial Indian Ocean sea surface temperature (SST) has also been emphasized to drive monsoon variability on interannual timescale (Li et al. 2001). On multidecadal to centennial or millennial timescale, the monsoon variability has been linked to changes in North Atlantic climate (Fleitmann et al. 2003; Gupta et al. 2003) and/or solar variability (Gupta et al. 2005). The Holocene record of the monsoon from the Oman Margin suggested that North Atlantic cold events and low solar activity are aligned with weak phases of the Indian summer monsoon (Gupta et al. 2003, 2005). To understand the orbital and suborbital variability in the climate of the equatorial Indian Ocean and its relation to glacial-interglacial cycles, we produced a 444 ka old proxy data from Ocean Drilling Program (ODP) Hole 716A, Maldives Ridge, equatorial Indian Ocean. Although the continuous and consistent sediment record of the monsoonal climate comes from the floor of the northwestern Arabian Sea (Gupta et al. 2003), ODP Hole 716A provides a continuous record of monsoon wind variability with moderate sedimentation rate free from terrigenous mixing.

Location, climatic and oceanographic setting ODP Hole 716A is located on the broad central plateau of the Maldives Ridge at a water depth of 533.3 m (04856.00 N; 73817.00 E), southeastern

Arabian Sea, equatorial Indian Ocean (Fig. 1). This hole lies in a flat terrain on a broad, shallow basin, and contains a virtually complete Cenozoic sedimentary record while remaining distant from any terrigenous influence throughout its development (Purdy & Bertram 1993). ODP Hole 716A is located just above the oxygen minimum zone (OMZ), has a well preserved sedimentary record, and is suitable for both faunal and stable isotope studies (Fig. 2). In the Maldives region, the winds are stronger from May –September or October and primary productivity reaches its maximum in September, that is, at the end of the SW monsoon, but is only moderate in comparison to high-productivity areas of the western Arabian Sea (Schulte et al. 1999). ODP Hole 716A lies under the influence of SW monsoon current although the wind intensity is not as strong as it is in the western Arabian Sea (Fig. 1). Present day bottom waters in the Maldivian region are well oxygenated owing to moderate productivity. The shift from the moist SW monsoon to the dry NE monsoon over the Indian subcontinent occurs during October and November. During this period, the NE winds contribute to the formation of the NE monsoon, which reaches the Maldives at the beginning of December and lasts until the end of March. However, the weather patterns of the Maldives do not always conform to the monsoon patterns of the Indian subcontinent. The winds blowing in a generally eastward direction during the summer are the strongest (Purdy & Bertram 1993). A deepening of the mixed layer is observed during the NE monsoon, but without an increase in primary productivity owing to the inflow of low salinity surface waters from the Bay of Bengal (Schulte et al. 1999). Another important climatic feature affecting the Indian monsoon is the Intertropical Convergence Zone (ITCZ). The ITCZ, also known as the meteorological or caloric equator, is a narrow latitudinal zone of wind convergence and precipitation. The shifts in the mean latitudinal position and structure of the ITCZ determine the onset, duration and termination of the rainy season in the tropics and subtropics on monthly to millennial timescales. Holocene palaeoclimatic evidence also suggests that such latitudinal shifts in the mean position of the ITCZ due to changes in insolation affect precipitation throughout the tropics thus bringing significant changes in the hydrological cycle (e.g. Fleitmann et al. 2007). In spring, the ITCZ migrates north of the equator across the Indian Ocean and reaches its northernmost position (approximately 208N) during boreal summer. In autumn, the ITCZ begins to retreat southward by the development of an intense high pressure system over Central Asia and reaches its southernmost position (approximately 108S) during boreal winter.

CLIMATE OF THE EQUATORIAL INDIAN OCEAN

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(a)

(b)

Fig. 1. Monthly mean wind field for (a) July (summer monsoon) (b) January (winter monsoon). The vectors give direction for the wind field and contours give magnitude of the pseudostress. The contour interval is 20 m2s22 (modified after Potemra et al. 1991). Also shown is the location of ODP Hole 716A.

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Fig. 2. Vertical profile of dissolved oxygen content along east– west transect at 048560 N (based on the data from GEOSECS 1983). OMZ, Oxygen Minimum Zone.

Materials and methods We sliced 9.92 m core from ODP Hole 716A at every centimetre interval. Samples were processed using the standard procedures with necessary precautions to avoid contamination following Gupta & Thomas (2003). Each sample was soaked in clean tap water overnight in a labelled beaker with half a spoon of baking soda, to aid rapid desegregation of the sample matrix. The soaked samples were washed with a jet of water over a 63 mm size sieve. After washing each sample, the sieve was stained with methylene blue solution and then washed with water thoroughly to avoid contamination of next sample. The washed samples were transferred to beakers, oven dried at a temperature of c. 50 8C and then transferred to labelled glass vials after drying. Approximately 300 specimens of benthic and planktic foraminifera from a suitable aliquot from .125 mm and .150 mm size fraction, respectively, were identified, counted and their percentages were calculated. The percentage of pteropods was calculated from a total population of c. 300 specimens of foraminifera from the same size fraction. The average age interval per sample at ODP Hole 716A is 498 years (ranging from 99 –2000 years) based on linear interpolation of six AMS 14 C calibrated dates (up to 30 536.65 calibrated years before the present) (unpublished) and two

nannofossil datums: First Appearance Datum (FAD) of Emiliania huxleyi (260 000 years at 7 m) and Last Appearance Datum (LAD) of Pseudoemiliania lacunosa (460 000 years at 14 m) (Backman et al. 1988). The ages of the faunal datums were updated to the timescale of Berggren et al. (1995). The AMS 14C dates were calibrated using CALIB 5.0.2 program (Stuiver et al. 2006). The ages for the interval from 444 000 to 150 000 years are based on oxygen isotope values of Globigerinoides ruber from ODP Hole 716A tuned with SPECMAP stacked d18O record (Imbrie et al. 1984) by wiggle matching using AnalySeries 2.0.3 (Paillard et al. 1996) (Table 1). We have plotted the percent and detrended time series of planktic foraminifera including mixedlayer species Globigerinoides ruber and Globigerinoides sacculifer, and thermocline species Neogloboquadrina dutertrei combined with the percent pteropods and characteristic benthic foraminifera Cymbaloporetta squammosa, Sphaeroidina bulloides and Uvigerina proboscidea from ODP Hole 716A (Fig. 3; supplementary material SUP18413). The detrending of the faunal time series was done by passing a fifth order best-fit polynomial and smoothed taking 5-point average (Fig. 3). Continuous Wavelet Transform (CWT) was performed on a benthic foraminifer Cymbaloporetta squammosa, planktic foraminifer Globigerinoides ruber and pteropods applying Daubechies wavelets

CLIMATE OF THE EQUATORIAL INDIAN OCEAN

Table 1. ODP Hole 716A ages and their source used in the present study Depth (mbsf) 0.19 0.32 0.51 0.70 0.90 1.07 5.54 6.72 7.00 7.56 8.19 9.71 14.00

Age (years bp) 4292.00 5616.00 7504.00 17 850.00 23 220.00 30 536.65 199 870.75 242 943.17 260 000.00 287 090.74 338 004.15 416 434.78 46 000 000.00

Source AMS 14C date* AMS 14C date* AMS 14C date* AMS 14C date* AMS 14C date* AMS 14C date* d18O age** d18O age** FAD of Emiliania huxleyi*** d18O age** d18O age** d18O age** LAD of Pseudoemiliania lacunosa***

*AMS 14C ages were calibrated with CALIB 5.0.2 (Stuiver et al. 2006). **Tuned with SPECMAP stacked age model (Imbrie et al. 1984). ***After Backman et al. (1988) updated to the timescale of Berggren et al. (1995).

using MATLAB 7.1 (Fig. 4). Prior to analysis, the values were interpolated at every 500 years applying Piecewise cubic Hermite interpolation (pchip) using Matlab 7.1. The data interpolation was consistent with the average temporal resolution of Cymbaloporetta squammosa, Globigerinoides ruber and pteropods data.

Results Both planktic and benthic foraminiferal time series show suborbital and orbital rhythms in various frequencies at ODP Hole 716A. Benthic foraminifer Cymbaloporetta squammosa shows an abrupt decrease and Sphaeroidina bulloides a significant increase at about 300 ka bp coinciding with the MidBrunhes Climatic Event (Fig. 3; Jansen et al. 1986; Kawamura et al. 2006). Uvigerina proboscidea shows a significant increase at c. 150 ka bp (Fig. 3). Pteropods show an increase whereas Gs. ruber shows a long-term decrease at c. 300 ka. Globigerinoides sacculifer shows a decrease at c. 400 ka and a major increase at c. 150 ka. Neogloboquadrina dutertrei follows somewhat similar trend as that of Gs. sacculifer (Fig. 3). These changes are in contrast to those from the northwestern Arabian Sea where marine sediments archive the record of SW monsoon-induced variability at various timescales (Emeis et al. 1995).

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The CWT of C. squammosa shows a shift from low frequency (40 ka) peak to high frequency peaks across the Mid-Brunhes Climatic Event (350 –250 ka) whereas Gs. ruber shows high frequency (suborbital) peaks during 444–250 ka (Fig. 4). The CWT of pteropods also shows a transition from low frequency (obliquity) peak to high frequency (precession) peak across the MidBrunhes Climatic Event and again over the past 50 ka. The high frequency peaks in benthic and planktic foraminifera could either be due to artefacts given the temporal resolution of the samples or solar-induced frequencies as observed in numerous records from the North Atlantic and the NW Arabian Sea (Mayewski et al. 1997; Gupta et al. 2003).

Planktic foraminifera and pteropods: their microhabitats Planktic foraminifera have been used extensively in palaeoclimatic reconstructions, especially during the Cenozoic because of their importance as indicators of surface water mass stratification since they live in specific layers of the world ocean (Berggren 1978; Jenkins 1993). Planktic foraminifera are sensitive to temperature, salinity, nutrient concentration, primary production and to a lesser extent oxygen content of the surface waters and may be helpful in reconstructing monsoon induced changes at ODP Hole 716A.

Mixed layer or shallow dwelling planktic foraminifera: Globigerinoides ruber and Globigerinoides sacculifer Globigerinoides ruber is a spinose planktic foraminifer living in the photic zone (top 50 m) of the water column in the tropical and subtropical areas with a SST of 14–30 8C (optimum 21–28 8C), optimum sea surface salinity (SSS) of 34.5– 36.0 psu (Be´ & Hutson 1977). It is a symbiontbearing species (with zooxanthellae), preferring oligotrophic regions with a deep mixed layer (Zheng et al. 2005) and is susceptible to dissolution (Berger 1971; Be´ 1977). Globigerinoides ruber most commonly occupies the warm mixed layer above the thermocline (Fairbanks et al. 1982) and shows maximum abundance in the top 20 m of the mixed layer in the early autumn when the thermocline begins to break down (Beveridge & Shackleton 1994). Globigerinoides sacculifer is also a shallow dwelling (surface depth habitat top 50 m) mixed layer, tropical to subtropical species, tolerating a SST range of c. 17 –30 8C (optimum 27 –30 8C), SSS

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Age (Ka)

Fig. 3. Raw per cent and detrended (5 point average) time series of planktic foraminifera (a) Globigerinoides sacculifer (b) Globigerinoides ruber and (c) Neogloboquadrina dutertrei combined with that of pteropods (d) and benthic foraminifera Cymbaloporetta squammosa (e) Sphaeroidina bulloides (f) and Uvigerina proboscidea (g) during the past c. 450 cal ka bp at ODP Hole 716A. Also marked are MIS stages in the top panel (after Imbrie et al. 1984).

CLIMATE OF THE EQUATORIAL INDIAN OCEAN 23

Fig. 4. Continuous Wavelet Transform (CWT) of benthic foraminifer Cymbaloporetta squammosa, planktic foraminifer Globigerinoides ruber and pteropods applying Daubechies wavelets. Prior to analysis, the values were interpolated at every 500 years applying Piecewise cubic Hermite interpolation (pchip).

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optimum of c. 34.9 psu and phosphate content of c. 0.5 mg l21 (Be´ & Hutson 1977). Globigerinoides ruber tolerates higher salinities whereas Gs. sacculifer prefers lower salinities (Bijma et al. 1990). Globigerinoides sacculifer is susceptible to dissolution and prefers low seasonal changes in SST and vertical temperature gradients and is not well suited to large seasonal salinity changes (Be´ 1977). Duplessy et al. (1980) have suggested that Gs. ruber secretes its shell within the mixed layer while Gs. sacculifer secretes its last chamber below the thermocline. Globigerinoides ruber (white) and Gs. sacculifer mirror regions of highest surface water temperatures and low primary production in the central and southern Arabian Sea (Ivanova et al. 2003).

Intermediate dwelling planktic foraminifer: Neogloboquadrina dutertrei Herbivorous foraminifer Neogloboquadrina dutertrei is a tropical to subtropical species with a SST range of 13 –30 8C (optimum 17– 25 8C), optimum SSS of 35.2 psu, phosphate content at 100 m around 0.6 mg l21 and living in top 100 m like G. bulloides (Be´ 1977; Be´ & Hutson 1977; Spooner et al. 2005). Neogloboquadrina dutertrei lives in active current systems, along continental margins (Cullen 1981) and in upwelling regions with chlorophyll maximum associated with the thermocline (Cannariato & Ravelo 1997). This species is abundant in high productivity areas (Marchant et al. 1999) with low salinity (Be´ 1977). Andreasen & Ravelo (1997) suggested that surface-dwelling (deep-dwelling) species increase (decrease) in abundance when the thermocline deepens. Modern observations reveal that different phases of upwelling in the Indian Ocean can be monitored using different species; N. dutertrei for the initial phase and G. bulloides for the final phase of upwelling (Kroon & Ganssen 1989). Neogloboquadrina dutertrei has been applied as an upwelling indicator in significantly weaker upwelling regions of the South and East China Seas where G. bulloides is not a dominant species (Jian et al. 2001).

Upper and intermediate water masses: the pteropods The population abundance of pteropods is influenced by the properties of the water column, and preservation is influenced by the bottom water. Owing to their larger size and mass, pteropod tests have a higher settling velocity that promotes deposition close to their habitat (Jasper & Deuser 1993). Preservation of pteropods declines with growing supply of organic carbon, oxygen

consumption and greater production of carbon dioxide. Well preserved pteropod shells are limited to the marginal seas with anti-estuarine circulation such as the Red Sea and the Mediterranean Sea (e.g. Almogi-Labin et al. 1991). In addition, pteropods accumulate in shallow parts of the open ocean, especially the tropical and subtropical oceanic regimes such as the Arabian Sea (e.g. Singh et al. 2001) and in environments which border carbonate platforms (Droxler et al. 1988). Upwelling regions from around the world recorded poor pteropod preservation (Berger 1978; Gerhardt & Henrich 2001). Numerous studies in the Arabian Sea have shown that during the interglacials and interstadials the SW monsoon is intense which leads to an increase in upwelling and bioproductivity off Somalia and Oman (e.g. Overpeck et al. 1996; Reichart et al. 1998; Gupta et al. 2003). Here, the main cause of dissolution of aragonite is the high input and remineralization of organic matter leading to high dissolved inorganic carbon (DIC) concentrations, which in turn lowers the pH of the intermediate (subsurface) waters (Millero et al. 1998). Also, higher export, burial and decomposition of organic matter within the sediments lower the pore water pH and contribute to dissolution of pteropods (Milliman et al. 1999). On the other hand, during cold intervals of reduced SW monsoon intensity, upwelling induced bioproductivity weakens and the preservation of pteropods increases owing to less corrosive bottom waters.

Benthic foraminifera: Cymbaloporetta squammosa, Sphaeroidina bulloides, Uvigerina proboscidea Cymbaloporetta squammosa has a well established microhabitat of low to intermediate oxygen and sustained flux of organic matter from high surface productivity (Sarkar & Gupta 2009; Sarkar et al. 2009). This species shows a good parallelism with a well-established upwelling indicator planktic foraminifer Globigerina bulloides and can be used as a proxy for wind strength at ODP Hole 716A (Sarkar & Gupta 2009). Sphaeroidina bulloides thrives in well-oxygenated water and is intolerant to oxygen depletion (Barmawidjaja et al. 1992). In the Indian Ocean, the association of this species with Cibicides wuellerstorfi and Anomalina globulosa indicates cool, active currents, low to intermediate organic flux, high seasonality and high oxygenation (Gupta & Thomas 2003). Uvigerina proboscidea is positively correlated with the organic carbon flux and negatively with the dissolved-oxygen concentration in the eastern Indian Ocean as higher primary productivity levels at the sea surface oxidize organic matter leaving the water

CLIMATE OF THE EQUATORIAL INDIAN OCEAN

oxygen-depleted at depth (Gupta & Srinivasan 1992). Peak abundances of U. proboscidea are inferred to represent times of high surface productivity related to intense trade winds during the SW Indian monsoon, causing widespread upwelling along equatorial divergence in the Indian Ocean (Gupta & Srinivasan 1992; Gupta & Thomas 1999).

Variability in equatorial wind intensity at ODP Hole 716A The population abundances of C. squammosa show major fluctuations between c. 450 and 300 ka with significantly higher values especially during warm intervals coinciding with markedly low pteropod population (Fig. 3). The population of this species abruptly decreases at c. 300 ka across MIS 8/9 and remains low thereafter indicating weakening of equatorial winds and low surface productivity over the past 300 ka. An abrupt decrease of C. squammosa population at c. 300 ka coincides with the Mid-Brunhes Climatic Event (MBCE), the beginning of aridity in the Indonesian and Australian region and a humid phase in equatorial East Africa (Jansen et al. 1986; Longmore & Heijnis 1999; Kawamura et al. 2006). Sphaeroidina bulloides (a low productivity species) shows a significant increase across MBCE and a decrease at c. 150 ka, following C. squammosa populations. Globigerinoides ruber and Gs. sacculifer, in general, show higher abundances during glacial stages, suggesting that during the glacial periods there was a deepening of the mixed layer and/or increased fresh water supply from the Bay of Bengal resulting from intense winter monsoon (Sarkar et al. 2000). The glacial changes at ODP Hole 716A suggest a reduced productivity and/or salinity in the Maldivian region. The pteropod population increases during glacial intervals, implying weaker winds and deepening of the OMZ (Fig. 3). The N. dutertrei percentages are moderate during these intervals. The decreased surface production and low organic flux to the sea floor during cold intervals led to less corrosive (perhaps better oxygenated) deep-sea conditions, enabling better preservation of pteropods at ODP Hole 716A. Similar pteropod preservation pattern was reported from core 905 off Somalia, core 137KA off Pakistan margin and core KL15 in the Gulf of Aden which show preservation spikes during stadials (von Rad et al. 1999; Almogi-Labin et al. 2000; Klo¨cker & Henrich 2006). An increase in Gs. sacculifer, and Uvigerina proboscidea coincides with a second phase of intense aridity in the Indonesia-Australian region (Longmore & Heijnis 1999; Kawamura et al. 2006). A major decrease in the intensity of westerlies in the equatorial Indian Ocean during the

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Mid-Brunhes Climatic Event coinciding with the beginning of an arid phase in the IndonesianAustralian region and a humid phase in equatorial East Africa signifies the role of equatorial Indian Ocean winds in the moisture budget of the Indonesian-Australian and African continents. The cause of decrease in wind intensity in the equatorial Indian Ocean across MIS 8/9 is still not fully understood. However, a weakening of Southern Oscillation phase linked to increased amplitude of winter insolation at equator may have driven changes in the Indian Ocean equatorial westerlies over the past 300 ka.

Conclusions Proxy records of equatorial Indian Ocean climate over the past 444 ka from the Maldives region (ODP Hole 716A) indicate significant changes in the equatorial wind intensity. The wind intensity shows a major increase during 450– 300 ka (MIS 12-9), leading to pronounced changes in the deepsea with deepening and shoaling of the oxygen minimum zone in the study area as can be observed in pteropod population at ODP Hole 716A. A major decrease in Cymbaloporetta squammosa and a significant increase in Sphaeroidina bulloides at c. 300 ka across MIS 8/9 suggest a weakening of equatorial winds in the Maldives region, which may have driven significant changes in distant places like the Indonesian-Australian and East African regions. A weak Southern Oscillation phase may have driven such a change. Ocean Drilling Program (ODP) is gratefully acknowledged for providing core samples for the present study (ODP sample request #17649E). This study was supported by DST, New Delhi (No. SR/S4/ES-46/2003) and CSIR, New Delhi through independent fellowship to SD and SS. Thoughtful reviews by Wolfgang Kuhnt are gratefully acknowledged. W.K. Mohanty, A. Routray, D. Pandit and S. Dhubia are thankfully acknowledged for their help in providing MATLAB 7.1 and performing CWT.

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Aeolian sediment evidence that global cooling has driven late Cenozoic stepwise aridification in central Asia H. LU*, X. WANG & L. LI School of Geographic and Oceanographic Sciences, Institute for Climate and Global Change Research, Nanjing University, Nanjing 210093, China *Corresponding author (e-mail: [email protected]) Abstract: It has been a long held view that uplift of the Tibetan Plateau dominated stepwise climatic drying in central Asia during the late Cenozoic. On the other hand, global cooling may also have forced Asian drying and the subsequent formation of aeolian deposits in north China. Until now, whether the Tibetan uplift or the global cooling has been the first-order driver controlling stepwise Asian drying has remained a contentious issue. In this study, we examine the thick aeolian silt deposit, which is regarded as a good archive of palaeoclimatic changes in central Asia and north China, in order to qualitatively reconstruct the drying process in Asia during the late Cenozoic. On the basis of our long-term field surveys, laboratory analyses and previous investigations, we have obtained time sequences of Asian drying from the early Miocene to late Pleistocene; we compare this newly reconstructed time series of Asian aridification with the time series of global cooling and Tibetan uplift to identify the first-order driver of stepwise Asian aridification. A good match between the drying and global cooling might indicate that global cooling was the most likely driver of stepwise drying in interior Asia. On the other hand, controversy regarding timing and amplitude of Tibetan uplift during the late Cenozoic suggests that the prevailing conclusion that Tibetan uplift forces Asian drying should be regarded as immature. A mechanism that global cooling drove the Asian drying is tentatively suggested.

There is an extensive dryland covered by the Gobi and deserts in central Asia which is an important dust source on the Earth (Fig. 1a). Dust emitted from these places is entrained and deposited in an area of regional and even hemispheric scale, influencing local and regional environments (Husar et al. 2001; Zhang et al. 2003; Jickells et al. 2005). Evolution and change of this dry environment has influenced hundreds of millions of people and even climate change beyond this region, and has attracted much investigation (e.g. Liu & Ding 1998). Most of the wind-blown silt deposits, namely the thick loess and the Red Clay (a loess-like silt deposit underlying the loess, and deposited in the Miocene and Pliocene) in north and central China, cover a time from the early Miocene to the Holocene, are deposited in the downwind area of the Gobi and deserts. These aeolian silt sequences are direct indicators of environmental changes in central Asia, and many palaeoclimatic reconstructions have been carried out using these aeolian silt sequences over the past two decades (Liu 1985; Kukla & An 1989; Zheng et al. 1992; Ding et al. 1998; Sun et al. 1998; Guo et al. 2002, 2008). However, long continuous palaeoclimatic records of these drylands during the late Cenozoic are incomplete, because the surface sediments in the Gobi and deserts are moveable and easily eroded.

In addition, the causal mechanism for the development of the dry environment has been hotly debated for many years without resolution. One main view is that the stepwise uplift of the Tibetan Plateau has directly driven stepwise Asian aridification during the late Cenozoic (Manabe & Terpstra 1974; Zhang 1981; Kutzbach et al. 1989; Ruddiman & Kutzbach 1989; Manabe & Broccoli 1990; Ruddiman 1997; Li 1999; Li & Fang 1999; An et al. 2001; Zhang et al. 2006; Fang et al. 2007; Sun et al. 2009). Other researchers suggest that global cooling controlled the stepwise drying in interior Asia, and that the growth of the Tibetan Plateau played a subordinate role (Zhou 1963; Chen et al. 1990; Wang & Gao 1990; Liu & Ding 1993; Guo et al. 1998, 2004; Dupont-Nivet et al. 2007; Jiang & Ding 2008; Jiang et al. 2008). Thus, more work is needed to clarify which hypothesis (global cooling or the Tibetan uplift) is the main driver of stepwise Asian drying in late Cenozoic. For this reason, we have reconstructed a time series of Asian drying during the late Cenozoic using wind-blown silt archives (the loess and the Red Clay) in north China. This newly obtained time series of Asian drying will then be compared with that of global cooling and with the uplift of the Tibetan Plateau to assess the possible forcing mechanism of Asian drying.

From: Clift, P. D., Tada, R. & Zheng, H. (eds) Monsoon Evolution and Tectonics –Climate Linkage in Asia. Geological Society, London, Special Publications, 342, 29– 44. DOI: 10.1144/SP342.4 0305-8719/10/$15.00 # The Geological Society of London 2010.

30 H. LU ET AL. Fig. 1. (a) The drylands and aeolian deposits in modern Asia. The dark area indicates the dryland in interior Asia, from which the dusts are emitted; the white-empty and empty arrows indicate the wind directions that transport dust to Chinese Loess Plateau and the dust storm track in spring in north Hemisphere, respectively. The inserted figure shows our filed survey routes, which provides the data for Table 1 and Figures 2, 3 and 4. (b) The index map of locations number 1287 that are listed in Table 1, and provide data sources for Figure 2.

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Fig. 1. (Continued).

Stepwise drying of central Asia during late Cenozoic Chinese loess and Red Clay as good indicators of aridification in Asia The wind-blown loess and Red Clay in north China are regarded as good indicators of dry environment

and of environmental changes in central Asia, because these aeolian sediments have a source in the Gobi and the deserts in interior Asia. Moreover, the northwesterly wind plays a key role of conveyer to link the dust source and sink, so that loess and Red Clay deposits preserved information of environmental changes in the dryland in central Asia (Liu 1985; Liu & Ding 1998; Rea et al. 1998; Husar

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et al. 2001; Zhang et al. 2003) (Fig. 1a, b). It has been shown that dry climates lead to extension of the Gobi and the deserts, strengthened physical weathering and wind erosion, produced more dust/silt for transport and deposition in downwind locations (Liu & Ding 1998). That caused an increase in the dust sedimentation rate, a coarsening the silt grain-size and an extension of aeolian silt cover in the downwind area at orbital and longer timescales (Liu 1985; Kukla & An 1989; Liu & Ding 1998; Guo et al. 2002; Stevens et al. 2007). The loess and the Red Clay are therefore sensitive to environmental changes in the Gobi and the deserts in central Asia. Conversely, reduced aridity of interior Asia reduces wind erosion, dust emission and transportation. Thus, the dust sedimentation rate becomes lower, dust particles are smaller and dust cover is reduced. As a result, changes in the loess –Red Clay sequences reveal climatic changes in a large region of the Gobi and the deserts, and can be used to reconstruct palaeoenvironmental changes in central Asia. Here, we use the dust sedimentation rate, grain-size distribution and cover area of the loess and the Red Clay as proxy indicators of the wet and dry changes in interior Asia at a tectonic timescale to reconstruct development of dry climates during the late Cenozoic (the data sources are listed in Table 1 and presented in Fig. 1b ).

Stepwise drying of interior Asia revealed by the aeolian deposits The existing investigations demonstrate that the loess and the Red Clay deposits in the Chinese Loess Plateau can be extended back to the early and middle Miocene (Guo et al. 2002, 2008; Lu et al. 2004a; Garzione et al. 2005; Fan et al. 2006; Qiao et al. 2006; Wang et al. 2006), clearly indicating that a dry environmental condition has occurred in central Asia since then. The Linxia fluviallacustrine deposit contains dust particles at the early Oligocene (Garzione et al. 2005), and may provide evidence for the earliest dust activity under arid environment in interior Asia. Commencement of the loess deposition at Qin’an site, dated at around 22 Ma (million years ago) (Fig. 2a; Table 1), probably indicates a sizable arid/semiarid land in central Asia, where the emitted dust was entrained and deposited in downwind regions, replacing the fluvial-lacustrine deposit of the Oligocene. The Miocene loess deposit at Qin’an also shows that there are two episodes of higher dust sedimentation rate and dust particle coarsening at around 15 –13 and 8– 7 Ma, respectively (Guo et al. 2002, 2008), revealing two dry phases in interior Asia. In addition, the aeolian silt deposit at Xining, around 700 km west of the Qin’an site, is dated back to

c. 14 Ma (Lu et al. 2004a; Wang et al. 2006), and probably demonstrates an extension of the windblown silt deposit and a strengthened drying in interior Asia during the middle Miocene (Fig. 2b; Table 1). This dusty event coincides with the dry climatic event revealed by a fluvial-lacustrine record in Linxia, east of Xining and west of Qin’an, and the Sikouzi site at Guyuan (Wang et al. 1999; Dettman et al. 2003; Garzione et al. 2005; Wang & Deng 2005; Fan et al. 2006; Jiang & Ding 2008; Jiang et al. 2008), showing it is at least a regional desiccation event. Further, an expansion of the wind-blown silt deposit in north China at around 8– 7 Ma has been well documented (Ding et al. 1998; Sun et al. 1998; Qiang et al. 2001) (Fig. 2b, c; Table 1). Investigation of grain-size distribution from the Red Clay along north–south and NW –SE transects for this period indicate a low-level wind playing a significant role on the dust transport, and that the dust provenance was located in the interior Asia (Miao et al. 2004; Wen et al. 2005). Many other investigations on palaeomagnetic stratigraphy and fossil assemblages have revealed this drying shift (e.g. Sun & Wang 2005) and the region of dust source was dry at this time evidenced by a recent direct dating (Sun et al. 2009). Following this drying shift, both dust sedimentation rate and grain size of the wind-blown silt deposit was significantly increased over the period of 3.6–2.6 Ma (Ding et al. 1997; An et al. 2001; Lu et al. 2004b; Wen et al. 2005) (Fig. 2d; Table 1), indicating a further significant drying event occurred in central Asia. This drying shift is also recorded by many other lines of geological evidence (Li 1999). At around 2.6 Ma, the Red Clay was replaced by the loess-palaeosol deposit which clearly indicates a further strengthened drying in Asia (Liu 1985; Kukla & An 1989; Liu & Ding 1998; Li 1999; Li & Fang 1999), showing at least a regional environment upheaval. The loess deposit was mantled over the Chinese Loess Plateau with coarser particles and higher sedimentation rates, compared to the underlying Red Clay deposit (Figs 2e, 3 & 4; Table 1). Meanwhile, contrast of the loess and the palaeosol of the aeolian silt sequence shows a strong cold-dry and wet-humid climatic oscillation, in pace with the glacialinterglacial changes in the Northern Hemisphere (Kukla & An 1989). Following this environmental shift, there are additional drying shifts revealed by the loess-palaeosol sediments at around 1.2 – 0.8 Ma (Fig. 2f; Table 1), which not only reveal a change from the obliquely-dominated cycles to the eccentrically-dominated cycles in this aeolian record, but an enhancement in contrast between the loess and palaeosol units (Lu et al. 1999, 2004b). After this shift, the loess deposit was extended to south of Qinling Mountains, which is

ASIAN ARIDIFICATION DRIVEN BY GLOBAL COOLING

33

Table 1. The investigated aeolian deposit sites in north China Number 1 2 3 4 5 6 7 8 9 10 11 12 13 14 15 16 17 18 19 20 21 22 23 24 25 26 27 28 29 30 31 32 33 34 35 36 37 38 39 40 41 42 43 44 45 46 47 48 49 50 51 52 53 54 55 56 57 58

Sites Qin’an Xining Bajiazui Baode Baoji Chaona Duanjiapo Fugu Jiaxian Jingchuan Lingtai Xunyi Zhaojiachuan Mazartag Sanju Guojialiang Heimugou Jingle Wubao Yecheng Ancun Baicaoyuan Baihe Baimapo Baishui Beiyuan Beiyuantou Caijiagou Caocun Caoxian Changwu Chenjiawo Chunhua Dadiwan Dadunling Dali Dunwashan Fuxian Gaolanshan Heshui Huangling Huanglong Huanxian Huining Jiezicun Jiuzhoutai Jiyuan Liujiapo Majiayuan Mangshan Mizhi Ningxian Panzishan Pingancun Pucheng Qinjiazhai Qishan Shangjiapo

Longitude (8)

Latitude (8)

References

Basal Age (Ma)

105.45 101.83 107.45 111.16 107.00 107.35 109.20 111.07 110.08 107.37 107.55 108.40 107.97 80.83 78.48 108.91 109.42 111.95 110.72 77.42 106.87 105.1 112.53 109.32 109.59 103.20 107.50 109.83 111.15 104.62 107.82 109.48 108.55 105.92 101.80 109.73 103.25 109.38 103.83 108.03 109.37 109.78 107.35 104.86 109.57 103.75 107.38 109.12 102.00 113.53 110.08 107.97 101.84 121.68 109.6 109.43 107.63 108.12

35.03 36.70 35.88 39.00 34.33 35.12 34.20 39.04 38.27 35.29 34.98 35.23 35.88 38.48 37.18 37.50 35.75 38.36 37.57 37.89 35.57 36.27 34.87 34.17 35.20 35.62 36.05 38.12 34.63 36.38 35.2 34.18 34.80 35.00 36.63 34.88 35.85 35.99 36.00 35.82 35.60 35.62 36.58 36.12 34.33 36.07 37.15 34.23 35.77 34.93 37.83 35.48 36.65 42.68 34.97 35.74 34.45 34.32

Guo et al. 2002 This study Sun et al. 1997 This study Evans et al. 1991 Ma et al. 2005 This study This study Qiang et al. 2001 Yang & Ding 2004 This study This study Sun et al. 1998 Sun et al. 2009 Sun & Liu 2006 This study This study This study This study Zheng et al. 2000 Sun et al. 2000 This study Sun et al. 2000 This study Sun et al. 2000 Sun et al. 2000 Sun et al. 2000 This study Sun et al. 2000 Sun et al. 2000 This study This study This study Sun et al. 2000 This study This study Sun et al. 2000 This study This study Sun et al. 2000 This study Sun et al. 2000 This study Sun et al. 2000 This study This study Sun et al. 2000 This study Sun et al. 2000 Sun et al. 2000 This study Sun et al. 2000 This study This study Sun et al. 2000 Sun et al. 2000 This study Sun et al. 2000

22  14 14  7.5

7.5  3.6

3.6  2.6

2.6  1.2

(Continued)

34

H. LU ET AL.

Table 1. Continued Number 59 60 61 62 63 64 65 66 67 68 69 70 71 72 73 74 75 76 77 78 79 80 81 82 83 84 85 86 87

Sites

Longitude (8)

Latitude (8)

Shimao Tuxiangdao Wuyishan Xiadongcun Xifeng Xueyuan Yanchang Yangguo Yichuan Yuanbao Chengdu Cele Chifeng Ganzi Kulungou Lijiagang Liuwan Naimanqi Shangbaichuan Shawan Taishanxincun Talede Xuancheng Xietongmen Yanziji Yining Yutian Zhenjiang Xigaze

110.00 101.73 103.22 110.67 107.70 106.97 109.93 109.52 110.17 103.17 104.90 80.72 118.82 99.98 121.25 118.93 110.13 121.16 110.05 85.61 118.71 83.25 118.85 88.24 118.81 81.51 81.68 119.68 88.90

37.92 36.58 35.80 36.10 35.70 36.92 36.66 34.35 36.06 35.63 30.52 37.17 42.32 31.63 42.62 32.92 34.13 43.22 34.07 44.34 32.15 43.52 30.90 29.46 32.14 44.00 36.85 32.22 29.32

the most important environment boundary of the North and South China, and, even to the Yangtze River catchments in southeastern China (Qiao et al. 2003; Lu et al. 2007). A further extension of loess cover in the last glacial period shows the aeolian silt deposit covers most parts of China (Fig. 2g; Table 1) (Liu 1985; Sun et al. 2000). A significant coarsening of the loess particles and a significant increase of the loess sedimentation rate during the late Pleistocene (Figs 3 & 4) indicates the most dry phase in Asia, manifested in various geological records (Liu 1985). Based on these aeolian silt records, the stepwise drying of interior Asia during the late Cenozoic are directly reconstructed (Figs 2–4).

Correlation of stepwise Asian drying and global cooling In the previous section, we drew a general picture of the stepwise Asian drying in the past c. 22 Ma (Figs 2 –4). If the stepwise drying in central Asia is compared to time series of global cooling (Shackleton et al. 1995; Zachos et al.

References This study This study Sun et al. 2000 Sun et al. 2000 This study Sun et al. 2000 This study This study Sun et al. 2000 This study Zhao et al. 2007 This study This study Sun et al. 2000 This study This study This study This study This study Fang et al. 2002a This study This study Qiao et al. 2003 Jin et al. 1996 This study This study Fang et al. 2002b This study Sun et al. 2007

Basal Age (Ma)

1.2  0.07

0.07  present

2001; Raymo et al. 2006), there is a striking correspondence at both orbital and longer timescales between the two time series, demonstrating a possible causal link. The appearance of a major ice-sheet in Antarctica at the Oligocene probably shows a link between global cooling and formation of the dry environment in Asia (Chen et al. 1990; Garzione et al. 2005; Barker et al. 2007). Formation of arid environments in the Oligocene is not easily interpreted in terms of Tibetan uplift forcing, because the height of the Tibetan Plateau at this time is not clear. Since there was only a small area of dust activity and the aeolian silt dust cover was not dominant in north China, the dry climate may not have been well developed and was probably weak in central Asia. By 24– 22 Ma, an increase in aeolian silt deposition may have coincided with the Mi-1 Glaciation (Zachos et al. 2001), which may show the cooling of the Earth was linked to the development of this Asian drying. Recent evidence shows that this wind-blown silt deposit covers an area of several hundred square kilometres, associated with suitable bed landforms (Guo et al. 2008), demonstrating an upheaval of the dry environment is a regional rather than a local one. The intensified

ASIAN ARIDIFICATION DRIVEN BY GLOBAL COOLING

drying event at c. 14 Ma may coincide with an important glacial expansion in Antarctica and the Arctic (Flower & Kennett 1993; Zachos et al. 2001). The remarkable drying in central Asia at around 8–7 Ma may be related to appearance of glaciers in the Arctic region (Moran et al. 2006), a rapid expansion of ice volume and a generally stronger cooling event on the Earth (Zachos et al. 2001). Further drying at 3.6–2.6 Ma in interior Asia may be associated with the rapid development of glaciers (Shackleton et al. 1995; Zachos et al. 2001). It is presumed that the reddish and fine Red Clay deposit was replaced by white and coarse loess at around 2.5 Ma and was associated with significant ice cover expansion in the Northern Hemisphere (Shackleton et al. 1984). Following this environmental shift, a drier climate dominated interior Asia from 1.2–0.8 Ma and probably coincided with expansion of the ice volume in the Northern Hemisphere after the Mid-Pleistocene Revolution (Shackleton et al. 1995; Zachos et al. 2001). A further expansion of the aeolian mantle in the last glacial period may be correlated with expansion of ice volume in the late Pleistocene and the Last Glacial Maximum (Bassinot et al. 1994; Shackleton et al. 1995). All the correlations between the dust deposition and the ice volume expansion may reveal a causal link between Asian drying and global cooling during the late Cenozoic (Fig. 5).

Decoupling of stepwise Asian drying and growth of the Tibetan Plateau The high relief of the Tibetan Plateau has certainly influenced Asian climate to some extent, for example, by blocking moisture from the Indian Ocean and South China Sea, and probably intensifying the Asian summer monsoon circulation and frequency and intensity of the cold surges (Asian winter monsoon) (Ruddiman 1997), which controls deflation, transportation and deposition of dust from central Asia (Liu & Ding 1998). However, timing and amplitude of tectonic uplift of the Tibetan Plateau is still debated, and timing and amplitude of uplift of the different parts are not well understood, so that any link between the phased uplift of the Tibetan Plateau and the stepwise drying in central Asia is yet to be determined. One concern is: by what date did the Tibetan Plateau reach a sufficient height to exert substantial influence over atmospheric circulation, and hence drive the Asian drying? The other consideration is whether each phased uplift have significantly changed the climate in interior Asia. There are still difficulties in quantifying the height and size of the Tibetan Plateau through time. From current knowledge, the Plateau (or an important part) may have reached a

35

significant altitude in the late Eocene and early Miocene (Rowley & Currie 2006; Clift 2006; Clift et al. 2006; Clift & Plumb 2008; Wang et al. 2008), the middle Miocene (Coleman & Hodges 1995; Spicer et al. 2003), the late Miocene (Ruddiman & Kutzbach 1989; Molnar et al. 1993; Ruddiman 1997; Molnar 2005), the Pliocene (Harrison et al. 1992; Zheng et al. 2000) or even that the plateau has grown significantly in the Quaternary (Li et al. 1979; Li 1999; Tapponnier et al. 2001). It is therefore not easy to appraise whether the Miocene or Pliocene uplift has substantially changed the environment in interior Asia, and further, we are not sure whether the Tibetan Plateau grew in a stepwise manner over the past c. 24 Ma. Even if many researchers suggested that the Tibetan Plateau uplift did force Asian drying (e.g. Ruddiman 1997), but to construct a direct linkage of the phased Asian drying and the Tibetan uplift at present time should not be easy; more evidence is needed to test this suggestion. On this basis, we suggest that the Tibetan uplift alone may not have been the direct driver of the stepwise Asian drying during the late Cenozoic.

Discussion Reliable dating of the aeolian silt deposit supports the proposal that global cooling drove stepwise drying in Asia Compared with lake and other geological records in Asia, there are many advantages to using the wind-blown silt deposit to reconstruct environmental change in interior Asia over a long timescale (million years). (1) It covers a range from several to tens of millions of years with continuity and has a time resolution of tens of thousands of years. This presents an ideal continental record, revealing regional rather than local environmental changes from the early Miocene to the Holocene. (2) The wind-blown silt deposit is well dated, so that a spatial stratigraphy correlation can be undertaken, and the timescale for this aeolian sequences is broadly accepted, with little controversy over its formation. (3) Because of its aeolian origin and relatively simple process of transport and deposition, developing a proxy index of the loess and the Red Clay for reconstructing environment changes is relatively straightforward. These features together offer a good opportunity to reconstruct long-term climate evolution in Asia in the late Cenozoic by the aeolian silt deposit in north China, and the new time series of the Asian drying obtained in this paper should be more complete and direct than that of the previous results, offering a unique

36

H. LU ET AL.

Fig. 2. Distribution of the aeolian deposit cover in late Cenozoic in north China, indicating development of dry climate in central Asia. (a) 22214 Ma (data from 2 sites); (b) 1427.5 Ma (data from 3 sites); (c) 7.523.6 Ma (data from 14 sites); (d) 3.622.6 Ma (data from 18 sites); (e) 2.621.2 Ma (data from 58 sites); (f) 1.220.07 Ma (data from 76 sites); (g) 0.07 Ma-present (data from 82 sites). These figures use data compiled from our long-term geological surveys and the references listed in Table 1.

opportunity to investigate the causal mechanism of the Asian drying. A good correlation between phased extension of the wind-blown silt cover and the global cooling during the Neogene does not unequivocally indicate

a causal linkage. However, when investigating both dust production/deposition and global climatic changes at a glacial-interglacial timescale, a correlation of the more dusty environment (namely a dry phase in Asia) and the glacial period (namely

ASIAN ARIDIFICATION DRIVEN BY GLOBAL COOLING

37

Fig. 2. (Continued).

a cool phase on the Earth) (Liu 1985; Kukla & An 1989; Liu & Ding 1993; Ding et al. 1995; Lu et al. 2004b) stimulates us to suggest a linkage between the global cooling and the Asian drying. This observation may reinforce our cooling-driven drying hypothesis. The view of the cooling-drying linkage may also be enhanced by consideration

of oscillators of the climatic changes in Asia during Pliocene and Pleistocene times. Wet-dry changes in Asia can be linked with global warmcold changes, but cannot be linked with changes in elevation of the Plateau height at an orbital timescale (indeed this assumption of elevation changes at orbital timescale does not exist). This may provide

38

H. LU ET AL.

Fig. 2. (Continued).

a new line of evidence that Asian drying was strongly tied to global cooling, but not with tectonic uplift. Our integrated aeolian records are in nearly one-to-one correlation with cooling phases on the Earth at the tectonic timescale (Fig. 5). This evidence supports our view of a coupling of the two systems.

A tentative model for global cooling driving Asian drying When the Earth’s climate is relatively cold, the temperature gradient between the poles and the tropics is increased, and the westerly jet in the north hemisphere is strengthened; at the same

ASIAN ARIDIFICATION DRIVEN BY GLOBAL COOLING

39

Fig. 2. (Continued).

Dust Sedimentation Rate (cm/ka)

time, cold surges in central Asia are strengthened and become frequent (Wang 2006). The strengthened cold wind brings less precipitation to interior Asia and entrains more dust to be deposited in downwind locations. Therefore, we can be confident that the global cooling has dried central Asia and consequently that the dry and windy climate has generated more dust particles. At the same time, a cold world causes sea level to fall and a consequent increas in continental land mass (the water source is further away from the interior of Asia, causing less

water to be brought into central Asia), thus an arid environment is formed. In the cold world, reduction of water circulation and a dustier environment (Yung et al. 1996) causes weaker moisture cycles in Asia, too. Therefore, a drier Asia was developed. Because of this reasoning, it is suggested that a cooling world is directly linked with the Asian aridification at both the orbital and tectonic timescales (Fig. 6). Development of dry climate in Asia is probably also linked to growth of the Tibetan Plateau, since

0

4

8

25

20

15

10

5

0

Age (Ma) Fig. 3. Changes in sedimentation rate of the aeolian deposits in late Cenozoic in north China, indicating development of the dry climate in central Asia. This figure is a synthesis of sedimentation rate data from Xining (Lu et al. 2004a) and Luonan (Lu et al. 2007) and references from Liu (1985; at Luochuan), Kukla & An (1989; at Xifeng), Ding et al. (1998; at Lingtai), Sun et al. (1998; at Lantian), Qiang et al. (2001; at Jiaxian), and Guo et al. (2002; at Qin’an).

40

H. LU ET AL.

Mean grain size (mm)

7

12

17

22 25

20

15

10

5

0

Age (Ma) Fig. 4. Coarsening of the aeolian silt deposits in the late Cenozoic in north China, indicating development of the dry climate in central Asia. Grain size data was averaged from all sampling levels with an interval of 20250 cm of the representative sections at Xining, Jiaxian, Lingtai, Lantian, Luochuan, Xifeng for this study. Qin’an site data is from Qiao et al. (2006). These aeolian silt sequences cover a time of past c. 22 Ma.

the high relief blocked moisture from the Indian Ocean and the South China Sea, strengthening the Asian monsoon circulations, as suggested by many researchers (e.g. Ruddiman 1997; Badgley et al. 2008). However, the connection between Tibetan uplift and Asian drying is far from conclusive because of poor constraints on the timing and amplitude of Tibetan uplift, whereas the connection between global cooling and Asian drying seems to be more robust. Uncertainties over the timing and amplitude of Tibetan growth means the tectonic forcing model should be further evaluated.

Challenges of the cooling forcing model We believe that a relatively complete picture of the Asian drying in the past c. 24 Ma has been

reconstructed from a study of the continuous windblown silt deposits; however, there are still weaknesses in this work that we should address in the future. One consideration is the completeness of the dust record. An area of around 1 000 000 km2 has been carefully investigated by us and other Chinese geologists in the past century, and the important aeolian sediment exposures of the late Cenozoic have been mostly reported (Fig. 1; Table 1); moreover, our field reconnaissance has covered most of the representative sites (Fig. 1b; Table 1). However, we believe that there are still some sites lost and that the current reconstruction of the dust cover is not complete. Further work is needed to find and investigate such sites (if they exist). On the other hand, surface erosion may also have caused loss of the early dust sediments, so

0 c. 22 Ma

δ18O(‰)

c. 14 Ma c. 8 Ma

c. 2.6 Ma

2

c. 1.0 Ma c. 0.07 Ma

Global Cooling 4 25

20

15

10

5

0

Age (Ma) Fig. 5. Coupling of the stepwise Asian drying and the global cooling during late Cenozoic. The d18O data is from Zachos et al. (2001). The arrows indicate drying events which occurred in central Asia, revealed by the aeolian silt records.

ASIAN ARIDIFICATION DRIVEN BY GLOBAL COOLING

Global cooling

Reduction of global hydrological cycle

Lowering sea level and increase in continentality

Strengthen cold-dry surges in Central Asia

Drying in Central Asia

Expansion of the wind-blown silt cover

Increase in dust sedimentation rate at downwind locations

41

We are grateful to Guo Zhengtang and Zheng Hongbo for their helpful suggestions; to Richard Bailey and Thomas Stevens for correcting the English; and, to Peter Clift, Ryuji Tada, the associate editor of the GSL book and two anonymous reviewers for significantly improving the paper. This research was supported by the National Natural Science Foundation of China (Grants no. 40930103, 30530050), the National Basic Research Program (2006FY110800) and the 111 project of Nanjing University.

References

Coarsening of particles at downwind locations

Fig. 6. A concept model to tentatively interpret the relationship between the Asian drying and the global cooling, please see the text for in details.

that what we have seen does not necessarily represent what existed. This problem may cause our current interpretation to be limited and overinterpreted. In addition, we need to examine how to quantitatively analyse the cooling-forced drying, because the timing correlation does not necessarily indicate a causal linkage. The last, and most difficult question, is that of the quantification of Tibetan uplift. This question is likely to remain unanswered for several decades, but is none the less important, especially in understanding Asian environment changes. Although there are many barriers to properly understanding the cause of Asian drying, our current integrated record may give a clue to answer this question.

Summary Our new data from the aeolian silt deposits in north China provide a record of stepwise drying of interior Asia. The good match between the degree of aridification and the stepwise cooling of the Earth suggests that global cooling may be the first-order driver of Asian drying during the late Cenozoic, through reorganization of atmospheric circulation and oceanographic current. The high relief of the Tibetan Plateau has certainly influenced the development of Asian drying, but it may have played a secondary role, superimposed on the Asian drying trend. Thus, we recommended that there is further evaluation of our proposal that Tibetan uplift caused drying in interior Asia, and also that investigations on the relationship between the Asian drying, global cooling and Tibetan uplift are still needed.

An, Z. S., Kutzbach, J. E., Prell, W. L. & Porter, S. C. 2001. Evolution of Asian monsoons and phased uplift of the Himalaya-Tibetan Plateau since Late Miocene times. Nature, 411, 62– 66. Badgley, C., Barry, J. C., Morgan, M. E., Nelson, S. V., Behrensmeyer, A. K., Cerling, T. E. & Pilbeam, D. 2008. Ecological changes in Miocene mammalian record show impact of prolonged climatic forcing. Proceedings of the National Academy of Sciences, 105, 12 145–12 149. Barker, P. F., Diekmann, B. & Escutia, C. 2007. Onset of Cenozoic Antarctic glaciation. Deep-Sea Research II, 54, 2293– 2307. Bassinot, F. C., Labeyrie, L. D., Vincent, E., Quidelleur, X., Shackleton, N. J. & Lancelot, Y. 1994. The astronomical theory of climate and the age of the Brunhes-Matuyama magnetic reversal. Earth and Planetary Science Letters, 126, 91– 108. Chen, M. Y., Hao, W. G., Yao, Y. Y. & Shao, M. Y. 1990. On the relationship between the Antarctic ice sheet and the eolian deposit in China. Quaternary Sciences, 3, 261– 271 (in Chinese). Clift, P. D. 2006. Controls on the erosion of Cenozoic Asia and the flux of clastic sediment to the ocean. Earth and Planetary Science Letters, 241, 571–590. Clift, P. D. & Plumb, R. A. 2008. The Asian Monsoon – Causes, History and Effects. Cambridge University Press, Cambridge. Clift, P. D., Blusztajn, J. & Nguyen, D. A. 2006. Large-scale drainage capture and surface uplift in Eastern Tibet before 24 Ma. Geophysical Research Letters, 33, L19403, doi: 10.1029/2006GL027772. Coleman, M. & Hodges, K. 1995. Evidence for Tibetan plateau uplift before 14 Myr ago from a new minimum age for east-west extension. Nature, 374, 49– 52. Dettman, D. L., Fang, X. M., Garzione, C. N. & Li, J. J. 2003. Uplift-driven climate change at 12 Ma: a long d18O record from the NE margin of the Tibetan plateau. Earth and Planetary Science Letters, 214, 267– 277. Ding, Z. L., Liu, T. S., Rutter, N. W., Yu, Z. W., Guo, Z. T. & Zhu, R. X. 1995. Ice-volume forcing of East Asian winter monsoon variations in the past 800 000 years. Quaternary Research, 44, 149–159. Ding, Z. L., Rutter, N. W. & Liu, T. S. 1997. The onset of extensive loess deposition around the G/M boundary in China and its palaeoclimatic implications. Quaternary International, 40, 53– 60. Ding, Z. L., Sun, J. M., Liu, T. S., Zhu, R. X., Yang, S. L. & Guo, B. 1998. Wind-blown origin of the Pliocene

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Desertification and dust emission history of the Tarim Basin and its relation to the uplift of northern Tibet RYUJI TADA1*, HONGBO ZHENG2, NAOMI SUGIURA1, YUKO ISOZAKI1, HITOSHI HASEGAWA1, YOUBIN SUN3, WENGANG YANG4, KE WANG4 & SHIN TOYODA5 1

Department of Earth and Planetary Science, Graduate School of Science, University of Tokyo, 7-3-1 Hongo, Tokyo 113-0033, Japan

2

School of Earth Science and Engineering, Nanjing University, Nanjing, 210093, China

3

Institute of Earth Environment, Chinese Academy of Science, 10 Fenghui South Road, Xi’an High-Tech Zone, Xi’an 710075, China

4

State Key Laboratory of Marine Geology, Tongji University, Shanghai 210092, China

5

Department of Applied Physics, Faculty of Science, Okayama University of Science, 1-1 Ridai, Okayama 700-0005, Japan *Corresponding author (e-mail: [email protected]) Abstract: The potential links between uplift of the Himalaya and Tibetan Plateau and desertification of inland Asia have been a long-considered problem in geology. Although a close link between the two has been suggested by theoretical climatic simulations, not enough geological data has existed to test the theory. Here, we conducted semi-quantitative field observations of a Neogene fluvial sequence at the Yecheng section on the southwestern margin of the Tarim Basin in order to confirm the origin and mode of deposition of the aeolian siltstone, determine the onset timing, evaluate quantitatively the temporal evolution of its deposition and its relationship to the tectonically driven surface uplift of NW Tibet. The results suggest a close link between the uplift of northwestern Tibet, alluvial fan formation, dust emission from Taklimakan Desert and the deposition of loess on the alluvial fans.

The possible link between uplift of the Himalaya and Tibetan Plateau (HTP) and evolution of the Asian monsoon has been a long-lasting and controversial problem of wide geological interest (e.g. Manabe & Terpstra 1974; Kutzbach et al. 1989; Ruddiman & Kutzbach 1989; Molnar & England 1990; Ruddiman 1997; An et al. 2001; Liu & Yin 2002; Kitoh 2004). The main reason for this controversy is our incomplete understanding of the timing and processes that control Asian monsoon evolution and the uplift of the HTP (An et al. 2001; Guo et al. 2002; Wang & Deng 2005; Harris 2006; Rowley & Curie 2006; Jiang & Ding 2008). In addition, insufficient spatial resolution and integrated time duration of palaeoclimatic model experiments using general circulation models (GCM), together with uncertainties that stem from changing boundary conditions outside the HTP region and the Asian monsoon system, make simulations of detailed parts of the Asian monsoon difficult. However, our understanding of the monsoon climate evolution has improved during the last decade because of the increasing number of

tectonic and palaeoclimatic studies, especially in central to northern parts of Tibet, as well as a drastic improvement in the performance of GCMs (e.g. Kitoh 2004; Rowley & Currie 2006; Royden et al. 2008; Wang et al. 2008). As a result, it has become possible to test the tectonic–climate linkage between several first-order features that are robust among many palaeoclimatic simulation experiments (Liu & Yin 2002; Abe et al. 2004; Kitoh 2004). One such feature is the link between the uplift of Tibet and the desertification of Central Asia in higher latitudes compared to other deserts that develop under subtropical high pressure zones (Liu & Yin 2002; Kitoh 2004). Deserts in inland East Asia are exceptional for their distribution at higher latitudes. Climatic simulations that include stepwise uplift of Tibet suggest that uplift results in the development of a wet climate in low latitude East Asia driven by intensification of the East Asian summer monsoon and the northward shift of desert areas to higher latitudes in the present position of the Taklimakan and Gobi deserts (40–508N). This shift reflects the rain shadow effect of Tibet (e.g. Kitoh 2004).

From: Clift, P. D., Tada, R. & Zheng, H. (eds) Monsoon Evolution and Tectonics –Climate Linkage in Asia. Geological Society, London, Special Publications, 342, 45– 65. DOI: 10.1144/SP342.5 0305-8719/10/$15.00 # The Geological Society of London 2010.

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Consequently, specifying the timing of formation and evolution of these deserts, and examining their relationship with the uplift of Tibet is an effective test for the proposed tectonic –climate linkages. Most previous studies of the timing of desert evolution in East Asia were based on studies of the loess sequence in the Chinese Loess Plateau (CLP) (Ding et al. 1998; Sun et al. 1998; Guo et al. 2001, 2002). Recent discovery of a Miocene loess sequence in the western CLP suggests the presence of deserts in inland East Asia at least as early as 23 Ma (Guo et al. 2002). Accumulation of loess continued and its extent and accumulation rates have increased in several steps since then (Ding et al. 2005). The most recent distinct increse occurred approximately between 3.6 and 3 Ma (Sun et al. 1998; An et al. 2001; Sun & An 2005). Similar increases in aeolian dust flux occurred at c. 3.6 Ma in the northern Pacific Ocean (Rea et al. 1998) and were interpreted to suggest expansion of deserts and/or enhancement of dust emission in East Asia. However, the exact location and extent of such deserts are not known because the source area of aeolian dust is not well constrained, although provenance studies using Nd and Pb isotopes suggest the Taklimakan as a possible source (Pettke et al. 2000). Aeolian deposits called ‘mountain loess’ are widely distributed in western China on the northern slopes of the Kunlun and Tian Shan Mountains that border the Tarim and Jungger Basins, respectively (e.g. Sun 2002). The mountain loess deposits are distributed at an altitude of up to 5300 m on the Kunlun Mountains. Its accumulation started around 1.1 Ma (Sun 2002). Because the earliest aeolian sands sampled in a borehole at the central part of the Taklimakan Desert are dated to be c. 1.1 Ma, Sun (2002) considered 1.1 Ma as the time of formation of the Taklimakan Desert. In contrast, Zheng et al. (2000) examined Miocene to Pliocene fluvial sequences of the Wuqia Group and the Artux and Xiyu formations in the southwestern Tarim Basin and reported the occurrence of probable aeolian deposits intercalated in the fluvial sequence. They also reported that the start of aeolian deposition is nearly synchronous with the start of deposition of conglomerate beds at c. 4.5 Ma, which these authors considered to reflect the onset of uplift of the western Kunlun Mountains in northwestern Tibet (Zheng et al. 2000, 2003, 2006). However, the origin (source), mode of deposition and depositional environment of these aeolian sediments were not examined in detail, and the nature of the linkage with the uplift of northwestern Tibet was speculative. The timing and mode of Tibetan uplift have been controversial for a long time, especially between Chinese scientists who have been considering this

problem dominantly from the northern side and other scientists, especially from Europe and North America, who have been tackling the problem from the Himalayan side (see Harris 2006 for review). However, a consensus seems to be emerging that Tibetan surface uplift occurred stepwise from the south to the north since the Early Cenozoic (Tapponnier et al. 2001; Rowley & Currie 2006; Royden et al. 2008; Wang et al. 2008). In brief, southern to central Tibet may have been uplifted to near present altitude by Eocene times, central to northern Tibet during late Oligocene–early Miocene, and northernmost and northeastern Tibet during the Plio-Pleistocene. The timing of Himalayan uplift seems less well constrained, but exhumation appears to have occurred during earliest Miocene (Godin et al. 2006), have been rapid during the middle Miocene (Clark et al. 2005) and probably during Plio-Pleistocene (Jain et al. 2000; Sakai et al. 2006). Knowledge of the more recent phases of the uplift, especially of northern Tibet, has also increased significantly during the last decade (Sun & Liu 2000). In particular, there seem to have been at least two pulses of uplift during the Plio-Pleistocene. The first pulse of uplift seems to have been around 5–3 Ma (e.g. Zheng et al. 2000) followed by a second pulse around 2–1 Ma (e.g. Sun & Liu 2000, 2006) in northwestern Tibet. A 5–3 Ma pulse of uplift has also been reported from northeastern Tibet (Fang et al. 2005). In this paper, we report on a detailed field mapping survey and observation of a ‘yellow siltstone’ facies of probable aeolian origin in the Artux and Xiyu formations at the Yecheng section in the southwestern part of the Tarim Basin in order to confirm the origin and mode of its deposition. In particular we consider the timing of onset and evaluate the temporal evolution of its deposition, as well as its relationship to the tectonic uplift of northwestern Tibet. We also conducted a preliminary study on the provenance of the ‘yellow siltstone’ to test the possible linkage between tectonic uplift of northwestern Tibet, its erosion and development of alluvial fans adjacent to northwestern Tibet, and the timing of dust emission and accumulation of loess on these alluvial fans. The results suggest a close link between the uplift of northwestern Tibet, alluvial fan formation, dust emission from Taklimakan Desert, and deposition of loess on the alluvial fans.

Geographical, climatological and geological setting of the studied area The Tarim Basin is a large (c. 560 000 km2) inland depression surrounded by the Kunlun Mountains to the south, the Pamir Plateau to the west, and the Tian

DESERTIFICATION OF THE TARIM BASIN

Shan to the north, whose average heights are in excess of 4000 m, with the highest peaks exceeding 7000 m (Fig. 1). Over 100 km wide alluvial fans develop on the foot of the surrounding mountains and fringe the margin of the Tarim Basin. These alluvial fans dip gently towards the basin floor with a slope angle of 2–58, and their surfaces form gravel deserts called gobi. Dry braided rivers (wadi) run across these alluvial fans and dry lakes (swamps) develop at the end of these braided rivers along the terminal margin of the alluvial fans. The floor of the Tarim Basin is occupied by the Taklimakan Desert, the second largest sand desert in the world. The elevation of the basin floor ranges from 800 –1300 m and dips slightly to the NE. Major rivers originating from the western Kunlun Mountains run towards the north or NE to merge into the Tarim River that flows eastwards along the northern margin of the basin and terminates near the northeastern end of the basin, where it forms large swamps and dry lakes (Fig. 1). There is no outlet river from the Tarim Basin. The Taklimakan Desert is one of the main source areas for dusts (mineral aerosol) to the Northern Hemisphere (Sun et al. 2001; Wang et al. 2004). It is also considered to be a major source for aeolian dust deposited on the CLP (e.g. Pettke et al. 2000). However, recent studies suggest that the

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sand desert is not the main source of the dust emission but rather show that the gobi on the alluvial fans, wadi running across the alluvial fans, and dry lakes near the margin of alluvial fans are more important sources (e.g. Derbyshire et al. 1998; Sun et al. 2001; Sun 2002). Along the southern margin of the Tarim Basin, fine-grained (silt size) pale, yellowish, grey aeolian deposits, called ‘mountain loess’, widely cover the upper slope of the alluvial fans and the lower part of the northern slope of the Kunlun Mountains (Sun 2002). Such mountain loess deposits are absent along the northern margin of the basin on the alluvial fans and the southern slope of the Tian Shan Mountains, most likely reflecting the surface wind system within the basin that consistently blows from the N – NE to the S –SW. The mountain loess is distributed at elevations in excess of 5000 m on the northern slope of the Kunlun Mountains (Derbyshire et al. 1998; Sun 2002) and is even contained in ice cores obtained from northern Tibet (Wu et al. 2004). In the present Tarim Basin, dust storms generally occur during late spring to early summer when the subtropical westerly jet jumps from its route to the south of Himalaya to a route to the north of Tibet (Liang & Wang 1998). This switch enhances the meandering of the subpolar westerly jet that in turn generates storms (Whitaker & Dole 1995).

Fig. 1. A simplified map of the Tarim Basin and surrounding area with the surface geographical conditions modified from Li et al. (1995). The locations of the studied site (Yecheng section) and Sanju section studied by Sun et al. (2008) are shown as stars.

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The change of the subtropical westerly jet to the north of Tibet also enhances the generation of a warm low-pressure cell over northern Tibet that allows a strong wind to blow towards Tibet, which then lifts up dusts emitted from the Tarim Basin to elevations in excess of 5000 m (Sun 2002). Dusts lifted above 5000 m are entrained in the subtropical westerly jet and transported long distances into the central Pacific Ocean, North America, or even to Greenland (Rea et al. 1998; Ruth et al. 2007). It is this special topographicalclimatological setting that makes the Taklimakan Desert unique. On the northern foot of the Kunlun Mountains in the southern margin of the Tarim Basin, a Neogene terrestrial sequence, deposited in terrestrial playa, proximal alluvial fan, and floodplain settings, is widely distributed (Zheng et al. 2000; Wang et al. 2006). Zheng et al. (2000) were the first to point out that the yellow siltstone beds intercalated in the sediments are of aeolian origin and that their accumulation started at c. 4.5 Ma at the Yecheng section in the southwestern part of the Tarim Basin (Zheng et al. 2000, 2003, 2006). Sun & Liu (2006) followed this line of study by examining the sequence at Sanju, c. 110 km to the east of the Yecheng section, and suggested that the onset of aeolian accumulation there was at 5.3 Ma. They argued that such aeolian deposition could provide significant evidence for the development of an arid climate and emergence of a desert as a dust source (Sun & Liu 2006; Sun et al. 2008). However, justification of the aeolian origin of yellow siltstones in these studies was based on the similarities in their grain size distributions, microstructures on grain surfaces, and geochemical characteristics compared with Quaternary loess deposits. No detailed examination of the field occurrence of the yellow siltstones was carried out.

Lithostratigraphy of the Yecheng section The lithostratigraphy and age model of the Neogene fluvial sequence at Yecheng was established by Zheng et al. (2000, 2006). The sequence is divided into the Miocene Wuqia Group, the Lower Pliocene Artux Formation, and the Upper Pliocene Xiyu Formation in ascending order (Fig. 2).

The Wuqia Group The Wuqia Group exposed at the Yecheng section is over 1700 m thick, and is characterized by alternations of red mudstone and a pale, brownish, fine-grained sandstone. The red mudstone shows some desiccation cracks and pedogenic features on its top, whereas the pale brownish sandstone

shows occasional current ripple structure at its top. Zheng et al. (2006), using sediment facies associations, interpreted the lower part of the Wuqia Group as having been deposited under meanderingstream conditions. The middle part was deposited in a braided-stream setting, and the upper part as deposited in a floodplain setting with episodic injections of coarse material into a muddy environment. Palaeocurrent directions in the Wuqia Group are consistently from east to west, suggesting the presence of a palaeo-river running from east to west along the northern margin of the Tibetan Plateau (Zheng et al. 2006). In this study we did not examine the Wuqia Group in detail and will not discuss further in this paper.

The Artux Formation The Artux Formation conformably overlies the Wuqia Group whose base is defined as the first appearance of pebble conglomerate (Zheng et al. 2006). The Artux Formation is composed of orange to yellowish grey fine-grained sandstone, orange, massive, very fine sandstone to siltstone, and granule to pebble conglomerates with minor thin-bedded red mudstone. The formation thickness within the Yecheng section is 810 m. The Artux Formation at Yecheng can be subdivided into three units (Fig. 2). The lower unit. The lower unit is c. 150 m thick, and is composed of two sedimentary cycles starting with conglomerate beds, fining up into decimetrescale alternations of reddish mudstone and yellowish grey fine sandstone, and then to decimetre- to metre-scale bedded orange to yellowish grey fine sandstone that thickens upward (Fig. 2). Conglomerates in the lower unit are generally subrounded to rounded and matrix-supported within a sandy matrix. Intercalations of red mudstone beds become rare to absent in the upper part of the lower unit, whereas decimetre to metre-scale alternations of yellowish grey, medium to fine sandstone and orange, very fine sandstone-siltstone become dominant up-section. Massive orange siltstone develops within the top few metres of each cycle. The middle unit. The middle unit is c. 230 m thick. A conglomerate bed, c. 1.2 m thick, develops at its base; most of the unit is characterized by decimetreto metre-scale alternations of yellowish-grey medium to fine sandstone and orange, very fine sandstone to siltstone (Fig. 2). Clasts within the conglomerates are generally subrounded to rounded and supported by a sandy matrix. An orange siltstone (over 5 m thick) is developed in the middle and the uppermost parts of this unit.

DESERTIFICATION OF THE TARIM BASIN

49

Fig. 2. A columnar section of the entire Neogene sequence (left) and that of the Artux Formation (right) at Yecheng.

The upper unit. The upper unit is c. 430 m thick, and is composed of six sedimentary cycles starting with conglomerate bed(s) changing upward into decimetre-scale alternations of reddish mudstone and yellowish grey fine sandstone (for the first, third, and fifth cycles), then into metre-scale cyclic repetitions of thick, orange to yellowish, very fine sandstone with decimetre-thick bedded yellowish grey fine sandstone and red mudstone (for the first, second, third and fifth cycles), and finally into thick (,50 m), massive, orange siltstone (Fig. 2). Intercalations of thin, red mudstone are more abundant in the upper unit, compared to the middle unit, especially in the lower part of each cycle. Intercalated beds of granule conglomerate become common from the fourth cycle and conglomerate beds become thicker upwards. These clast-supported conglomerates comprise subangular to subrounded components with a sandy or silty matrix.

Depositional environments. Zheng et al. (2006) interpreted the lower part of the lower unit (the lowermost 50 m that corresponds to the first cycle) as having been deposited in a floodplain setting, based on sedimentary facies associations. The interval corresponding to the main part of the lower unit and the middle unit was deposited in a distal fan setting, based on the occurrence of medium to coarse cross-bedded sandstone, interbedded with granule-grade conglomerate beds and very minor red mudstone. The upper unit was deposited in a lower to middle fan setting based on the dominance of bedded, fine-grained sandstones with planar laminations and homogeneous siltstone and lack of current indicative sedimentary structures, such as channels and ripples. Our observation basically confirms this interpretation. The dominance of red mudstone beds that occasionally show desiccation cracks alternating with thin greyish sandstone

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R. TADA ET AL.

beds in the lower part of the lower unit is consistent with a flood plain setting. Decimetre-thick, yellowish-grey, medium to fine sandstone beds with occasional, very thin (,1 cm), red, muddy drapes in the middle unit may represent flood events in a distal fan setting. In contrast, the occurrence of decimetre-thick, red mudstone beds, with thin sandstone intercalations in the lower part of the first cycle in the upper unit is similar to that in the lower part of the lower unit, and may represent recurrence of a floodplain setting. In summary, the lower and middle units plus the first cycle of the upper unit of the Artux Formation, are considered to represent sedimentation in flood plain to distal fan settings, whereas the main part of the upper unit (from the second cycle) of the Artux Formation represents a lower to middle alluvial fan setting. Zheng et al. (2006) also suggest that the massive, orange siltstone beds are of aeolian origin, based on their massive appearance, grain size, good sorting, and carbonate cementation. Such massive, orange siltstones of several metres to 50 m thickness occur in the uppermost part of each cycle (Fig. 3a). They are generally homogeneous and structureless, but have a few thin (,1 cm) reddish clayey stripes, probably representing deposition from overbank floods on the lower to middle part of an alluvial fan (Fig. 3b). Decimetre to metre-thick, orange, very fine sandstone to siltstone of similar appearance is also intercalated in the middle part of each cycle. These beds become common to dominant in the upper unit. Such orange, very fine-grained sandstone to siltstone might also be of aeolian origin. However, reworking of aeolian deposits by river flow is also possible, considering their occurrence as decimetre- to metre-scale rhythmical beds.

The Xiyu Formation The Xiyu Formation conformably overlies the Artux Formation with an erosional surface and its base is defined as the first appearance of thick pebble to boulder conglomerate. The Xiyu Formation is dominantly composed of pebble to boulder-sized conglomerate layers with occasional intercalation of siltstone layers and lenses (Zheng et al. 2006). The Xiyu Formation at Yecheng is c. 2100 m thick, and is subdivided into two units based on slight differences in lithological characteristics. The lower unit. The lower unit is logged as being 1040 m thick, and is characterized by thick conglomerate layers with intercalations of siltstone layers (,3 m) that are less common than those in the upper unit. Here we use the term ‘layer’ instead of ‘bed’ because some of the layers are amalgamated

Fig. 3. Photographs showing the field occurrences of aeolian sediments in the Artux Formation at Yecheng section. (a) A thick massive orange siltstone in the upper unit of the Artux Formation. (b) A thick homogeneous orange siltstone with red muddy stripes. The pen is c. 13 cm in length.

beds without any obvious intercalations. Conglomerates of the lower unit are subrounded to subangular, poorly-sorted, clast-supported, and generally structureless, but with occasional upward coarsening. The maximum diameter of clasts is generally between 5 and 30 cm. Although many of the conglomerate layers are structureless, faint parallel and cross stratifications are observed in the upper part of this unit (c. 7002800 m above the base of the Xiyu Formation) suggesting deposition under stream influence. Thickness of the conglomerate layers is commonly between 5 and 25 m, with an average of 17 m, but the maximum thickness exceeds 70 m. Orange siltstone lenses and lenticular beds commonly occur within conglomerates of the lower 300 m of this unit (Fig. 4a). More than 300 m above the base, the siltstone becomes yellower and occurs as layers rather than lenses (Fig. 4b). Yellow siltstone layers in the lower unit are less than 2 m thick, with an average thickness of 0.9 m. Orange siltstone lenses and yellow siltstone

DESERTIFICATION OF THE TARIM BASIN

51

Fig. 4. Photographs showing the field occurrences of aeolian sediments in the Xiyu Formation at Yecheng section. (a) A lens of orange siltstone in the basal part of the Xiyu Formation. (b) A decimetre-scale layer of yellow siltstone in the lower unit of the Xiyu Formation. (c) A yellow siltstone layer covering an irregular surface of the underlying conglomerate in the lower unit of the Xiyu Formation. (d) A 7.2 m thick yellow siltstone layer in the upper unit of the Xiyu Formation (at 1343 m above the base of the Xiyu Formation). (e) An intercalation of a thin red mudstone bed in a yellow siltstone layer in the upper unit of the Xiyu Formation. (f) A whitish grey cross-bedded fine sandstone in the uppermost part of the upper unit of the Xiyu Formation.

layers of the lower unit generally cover either irregular or flat surfaces of the underlying conglomerates, with sharp contacts in the lower to middle part, although gradual contacts become more

common in the upper part of this unit. The tops of orange siltstone lenses and yellow siltstone layers are generally eroded by overlying conglomerates forming irregular erosional contacts (Fig. 4c).

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Orange siltstone lenses and yellow siltstone layers are either homogeneous or decimetre-thick bedded, with occasional thin intercalations of red mudstone and granule-grade conglomerate. Faint laminations are occasionally found especially in the upper part of yellow silt layers in the upper half of this unit. The upper unit. The upper unit is c. 1160 m thick, and its base is arbitrarily defined by the first appearance of thick (.3 m) yellow siltstone layers at 1041 m above the base of the Xiyu Formation. The upper unit is characterized by more frequent intercalations of yellow siltstone layers than occur in the lower unit. The thickness of yellow siltstone layers becomes more variable, with a maximum thickness of 7.2 m (Fig. 4d). The upper unit is also characterized by a thicker and more variable grade conglomerate. Clasts in the conglomerates of the upper unit are subangular to angular, poorly-sorted, clast-supported, and either massive or faintly stratified. The maximum diameter of the clasts significantly increases from 70 cm to .1 m higher than 1185 m above the base of the Xiyu Formation. Although many of the conglomerate layers are structureless, faint parallel stratification are observed in the lower part of this unit (c. 1250 to 1420 m above the base of the Xiyu Formation) suggesting deposition under the influence of stream flow. The average thickness of a conglomerate layer is c. 14 m, but the maximum thickness exceeds 50 m. Yellow siltstone layers are variable in thickness, ranging from 0.5–7.2 m, with an average of 1.9 m. On large cliffs exceeding 100 m height and 500 m in width on which the Xiyu Formation is exposed, yellow siltstone layers thin towards the south and sometimes pinch out. Yellow siltstone layers of the upper unit generally cover flat surfaces of the underlying conglomerates with gradual contacts, whereas the tops of yellow siltstone layers are generally eroded by overlying conglomerates. Yellow siltstone layers are either homogeneous or decimetre-thick beds, with thin intercalations of granule grade conglomerate and reddish mudstone (Fig. 4e). Faint lamination is occasionally found, especially in the upper part of yellow siltstone layers. An c. 6 m thick, whitish grey, cross-bedded fine to medium sandstone with black grey scoriaceous coarse sandstone thin beds occurs near the top of the section (Fig. 4f). The thickness of one cross-bed set was observed to be less than 1 m and cross beds dip to the north at an angle of c. 208. Relatively low-angle and thin cross-bed sets argue against an aeolian origin and suggest a fluvial origin (Hunter 1977; Brookfield 1992). It is possible that this whitish grey crossbedded sandstone represents erosion following a volcanic event in northwestern Tibet that supplied

significant amount of pyroclastic debris to the southwestern margin of the Tarim Basin. Provenance of conglomerate. Zheng et al. (2006) examined the composition of conglomerates from the Artux and Xiyu formations in detail. According to these authors, the conglomerates are polymictic and composed of sedimentary clasts (including purplish mudstone, pink and grey limestone, greywacke, and sandstone), low-grade metamorphic clasts (such as greenschist, greenish metasiltstone, and marble), higher grade metamorphic clasts (e.g. gneiss), volcanic clasts (both silicic and mafic), and plutonic clasts (e.g. granite, dolerite, diorite and gabrro) (see Zheng et al. 2006 for details). Zheng et al. (2006) further examined temporal changes in the composition of conglomerate, and found that sedimentary clasts are common throughout the Xiyu Formation, whereas plutonic clasts only appear above 640 m from the base of the Xiyu Formation. Plutonic and highgrade metamorphic clasts become common above the 100021200 m level, close to the base of the upper unit.

Age model and sedimentation rates for the Artux and Xiyu formations A detailed chronology for the upper part of the Neogene sequence, the Artux and Xiyu formations, at Yecheng was established by Zheng et al. (2000) based on high-resolution magnetostratigraphy. We adopted this chronology and projected their polarity reversal boundaries onto our columnar section (Table 1). We did not use the bottom of Chron C2r1n as an age control because it is too close to the top of C2r1n and the estimated error for sedimentation rate in the interval corresponding to C2r1n is too large. Constant sedimentation rates have been assumed between the reversal boundaries. Based on the age model, linear sedimentation rates (LSRs) for the Artux Formation are between 579 and 1370 m/Ma with the highest LSR in the upper part of the middle unit. LSRs for the lower unit of the Xiyu Formation are very high, ranging between c. 1530 and 3020 m/Ma whereas the LSRs for the upper unit are lower between 386 and 1510 m/Ma (Fig. 5).

Tilting history of the Yecheng Section Throughout the entire sequence exposed at Yecheng, the strata have relatively stable strikes between N668W and N888W with .90% being within N818 + 78W. Dips are consistently to the north, and steepen upward slightly from the base of the exposed sequence (c. 8 Ma; Zheng et al.

Table 1. Stratigraphic positions of polarity reversals in the Artux and Xiyu formations, and equations for age estimations Corresponding polarity chron

0 1 2 3 4 5 6 7 8 9 10 11

C2n C2r1n C2An1n C2An1n C2An2n C2An2n C2An3n C2An3n C3n1n C3n1n C3n2n C3n2n

Top Age (Ma)

Bottom Age (Ma)

Duration (m.a.)

Top level (m)

Bottom level (m)

Interval thickness (m)

LSR (m/m.a.)

Age estimate equations

.1.77 1.95 2.14 2.58 3.04 3.11 3.22 3.33 3.58 4.18 4.29 4.48

1.95 2.14 2.58 3.04 3.11 3.22 3.33 3.58 4.18 4.29 4.48 4.62

,0.18 0.19 0.44 0.46 0.07 0.11 0.11 0.25 0.6 0.11 0.19 0.14

0 272 455 625 1043 1253 1436 1768 2150 2529 2680 2790

272 455 625 1043 1253 1436 1768 2150 2529 2680 2790 2921

272 183 170 418 210 183 332 382 379 151 110 131

.1511.1 963.2 386.4 908.7 3000.0 1663.6 3018.2 1528.0 631.7 1372.7 578.9 935.7

Age ¼ 1.77 þ 0.000662*d Age ¼ 1.95 þ 0.001038*(d - 272) Age ¼ 2.14 þ 0.002588*(d - 455) Age ¼ 2.58 þ 0.001100*(d - 625) Age ¼ 3.04 þ 0.000333*(d - 1043) Age ¼ 3.11 þ 0.000601*(d - 1253) Age ¼ 3.22 þ 0.000331*(d - 1436) Age ¼ 3.33 þ 0.000654*(d - 1768) Age ¼ 3.58 þ 0.001583*(d - 2150) Age ¼ 4.18 þ 0.000728*(d - 2529) Age ¼ 4.29 þ 0.001727*(d - 2680) Age ¼ 4.48 þ 0.001069*(d - 2790)

DESERTIFICATION OF THE TARIM BASIN

Zone

Positions of polarity reversals are after Zheng et al. (2000) and ages for polarity reversals are after Cande & Kent (1995). LSR, linear sedimentation rate.

53

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R. TADA ET AL.

1.5

2

2.5

Age (Ma) 3 3.5

4

4.5

1500

Upper Unit

Xiyu Formation

1000

Lower Unit

500

2500

3000

Artux Fm.

2000 Lw. Mid. Upper U.

Depth (m) from the top of the Xiyu Fm.

0

Wuqia Gr.

Fig. 5. A diagram showing the relationship between the age versus depth (from the top of the Xiyu Formation) of the strata in the Artux and Xiyu formations at Yecheng section.

2009) to the upper unit of the Artux Formation (c. 4.0 Ma), then decrease gradually towards the top of the Xiyu Formation (c. 1.8 Ma). Observation of the large exposure of the Xiyu Formation from a distance revealed that strata thin and/or pinch out towards the south without folding, suggesting syntectonic sedimentation [growth strata of Sun et al. (2008)] (Fig. 6). Consequently, it has been possible to reconstruct the tilting history by examining temporal changes in dip angles (Fig. 7). We reconstructed changes of the dip angle with the age of strata at Yecheng. As can be seen in Figure 7 that the dip angle increases slowly from 628 at c. 8 Ma to 808 at 4.0 Ma with the average rate of 4.58/Ma. At c. 4.0 Ma, a decrease in the dip angle starts, and continues to decrease with an average rate of 298/Ma until 3.2 Ma. The average rate of the dip angle decrease suddenly increases to 1208/Ma at c. 3.2 Ma and continues until 3.0 Ma, during which time interval the dip decreases from c. 578 to 338. The average rate of dip angle decrease falls to 188/Ma by c. 3.0 Ma and continues until 2.6 Ma. Dip angles remain at a nearly constant value of 268 from c. 2.6 Ma to 2.2 Ma. A decrease in dip angle resumes at c. 2.2 Ma at a rate of 338/Ma until at least 1.8 Ma when the dip angle is 138. If we extrapolate this decreasing rate to the dip angle of 3.58, an average original dip of alluvial fan deposits, we obtain an age of c. 1.5 Ma. It is uncertain whether the slow increase in dip angle from 8 Ma to 4.0 Ma is because of syndepositional tilting or not. In contrast, the decrease in the dip angle since 4.0 Ma is best explained by syndepositional tilting to the north (Zheng et al.

2000). Thus, our result suggests an onset of northward tilting at 4.0 Ma. There are two phases of northward tilting. The first phase of tilting was from 4.0 Ma to 2.6 Ma with a distinct maximum in tilting rate between 3.2 Ma and 3.0 Ma. Cessation of tilting occurred between 2.6 Ma and 2.2 Ma, and the second phase of tilting started from 2.2 Ma and probably lasted until c. 1.5 Ma. Similar syntectonic sedimentation was recently reported from Sanju where the tilting started at 5.3 Ma according to Sun et al. (2008), although detailed data to constrain the onset timing of tilting were not presented. Because there is an angular unconformity between the Xiyu Formation and the Pleistocene Wusu Formation (Zheng et al. 2000; Sun et al. 2008), and the Wusu Formation is nearly horizontal, we infer that there must have been another phase of tectonic activity that was not associated with tilting, but was rather characterized by simple uplift, approximately between 1.5 Ma and 1 Ma.

Linear sedimentation rates for aeolian siltstone and alluvial fan conglomerate In this study, we estimated LSRs for aeolian dusts and conglomerate, respectively, and reconstructed their changes from 4.6 Ma to 1.8 Ma at Yecheng. We measured thicknesses of yellow siltstone layers in the Xiyu Formation at the outcrops where their thicknesses exceed c. 1 m. We also measured the thicknesses of yellow siltstone layers on photographs of cliffs where they are between c. 0.3 and 1.0 m. Yellow siltstone layers and lenticular beds of orange siltstone thinner than c. 0.3 m were not measured because they were difficult to identify on photographs and are generally not continuous laterally. Thicknesses of conglomerate beds in the Xiyu Formation were either measured at outcrop or calculated from the route map. As for the Artux Formation, we measured thicknesses of thick (.1 m) homogeneous, orange siltstone layers either at outcrop or on photographs. We did not measure the thinner orange siltstone to very fine sandstone layers that alternate with yellowish grey sandstones and/or red mudstone layers because we are not confident about their possible aeolian origin. As a result our estimate of aeolian dust LSRs should be treated as an underestimate, especially in the middle unit of the Artux Formation, although this does not affect the main conclusion of this paper. Thicknesses of conglomerate beds in the Artux Formation were measured at outcrop. Errors in the thickness estimates are ,10% for yellow siltstone of the Xiyu Formation and conglomerate of the Xiyu and Artux formations because they are based on measurements on photographs of cliffs of which the width was measured directly,

DESERTIFICATION OF THE TARIM BASIN

Fig. 6. A distant view of a large cliff of the upper unit of the Xiyu Formation at Yecheng section. The dip of stratum gradually shallows upward without any folding, suggesting tilting occurred concurrent with deposition (growth strata). Width of the entire view of the photo is c. 3 km.

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Artux Formation

Xiyu Formation Upper

Dip (°)

80

Lower

Upper

Mid

L

Wuqia Gr.

56

(a)

60 40 20

Yel Slt LSR (m/Ma)

0 300

(b)

200

100 0

Maximum Average Minimum

Cgl. LSR (m/Ma)

3000 (c)

2000 1000 0 2

2.5

3 3.5 Age (Ma)

4

4.5

Fig. 7. Diagrams showing the changes in: (a) the dip angle of the strata; (b) LSR of orange and yellow (aeolian) siltstone; and (c) LSR of conglomerate with the age in the Artux and Xiyu formations. The error of dip angle represents an uncertainty of measurement. Additional uncertainty of +1.58 for the original dips of strata should be taken into account when reconstructing the tilting history.

whereas the error for the orange siltstone of the Artux Formation may be as large as 30% because the exposures are near the top of cliffs and only photographs are available. Figure 7b shows the changes in LSRs of yellow and orange siltstones of aeolian origin over time. The LSR of aeolian siltstone, which is negligibly low before 4.5 Ma, started to increase from that time and reached a maximum value of 250 m/Ma

between 4.2 and 3.6 Ma. Although the LSR of aeolian siltstone in Figure 7b suggests a significant increase after 4.2 Ma, this is an artefact caused by averaging the rate between 3.6 and 4.2 Ma. Our field observations suggest that orange siltstone to very fine sandstone occurs more commonly after 4.0 Ma. LSR decreases gradually from 3.6 Ma to a minimum value of 33 m/Ma between 2.6 Ma and 2.1 Ma. The LSR gradually increased from 2.1 Ma

DESERTIFICATION OF THE TARIM BASIN

to 217 m/Ma at 1.8 Ma. Figure 7c shows changes in LSR for conglomerate. Although deposition of conglomerate started at c. 4.6 Ma, its LSR was very low until 3.6 Ma. The LSR of conglomerate increased rapidly to c. 1290 m/Ma at 3.6 Ma and reached maximum values of 1590 to 2870 m/Ma between 3.6 Ma and 3.0 Ma, before decreasing to 350 m/ Ma between 2.6 and 2.1 Ma. Subsequently LSR gradually increased to 1260 m/Ma towards 1.8 Ma. This study revealed that the LSR of aeolian siltstone was as high as 250 m/Ma. This value is comparable or even larger than the LSRs for the Quaternary loess sequence in the CLP (e.g. Sun et al. 2006) and in mountain loess on the southern margin of the Tarim Basin (Fang et al. 2002), suggesting significant dust emission from the Tarim Basin. The LSRs of aeolian siltstones were lower in the Xiyu Formation compared to the upper unit of the Artux Formation, which may have been due partly to erosion during conglomerate deposition.

Grain size analysis of the aeolian deposits of the Artux and Xiyu formations The field observational results described above suggest an aeolian origin for the thick homogeneous orange siltstone of the Artux Formation and orange-yellow siltstone lenses and yellow siltstone layers of the Xiyu Formation. To extract further information on their depositional environment, we conducted grain size analysis of 41 samples of aeolian siltstones taken from the Artux and Xiyu formations. We also analysed grain size for 38 samples of the conglomerate matrix of the Artux and Xiyu formations in order to compare with aeolian siltstones from similar horizons. To isolate the detrital component, samples were crushed to granule size. They were treated with 20% CH3COOH to remove carbonate, then with 0.3 mol/l C6H5O7Na3.2H2O and 1 mol/l NaHCO3 with 8:1 ratio plus 3 g of Na2S2O4 to remove ferric and manganese oxides, and finally with 10% H2O2 to remove organic matter (Rea & Janecek 1981). We then removed the ,4 mm fraction, a minor component of the bulk samples, from the chemically-treated samples by the pipette method to avoid interference from flaky clay particles on laser scattering. These pre-treated samples were analysed with a Horiba LA-920 laser-diffractionscattering grain size analyser equipped with tungsten and He –Ne laser light sources. Approximately 120 mg of the pre-treated samples was suspended in 0.2% of sodium pyrophosphate (to disperse grains) and put into the analyser. Measurement was conducted over a diameter range of 0.0222000 mm with grain size class of every Dlog10F ¼ 0.06 (Nagashima et al. 2007). Reproducibility is within

57

+0.3 mm for median diameter and +0.012 for sorting. Nearly all of the samples show either unimodal distributions with a shoulder or bimodal distributions. We used Igor-pro software to de-convolute the grain size distribution into multiple log-normal distributions. Two log-normal distributions are generally enough to explain the original grain size distributions in the case of samples analysed in this study, with finer median diameter (¼ modal position) at around 10 mm and coarser median diameter at around 70 mm. For the finer and coarser log-normal distributions, we calculated median diameter, sorting, and content of each log-normal distribution. The median diameter of the finer distribution of orange and yellow siltstone ranges from 7– 17 mm, with an average of 11.2 mm. The coarser distribution ranges from 342102 mm, with the average of 72.4 mm. Figure 8 shows the changes in median diameter of the finer and coarser distributions of the aeolian (orange and yellow) siltstones in the Artux and Xiyu formations. Median diameter of the finer distribution was relatively constant at c. 9 mm in the main part of the Artux Formation (before 4.0 Ma), then increased gradually from the upper part of the Artux Formation towards the top of the Xiyu Formation (c. 1.8 Ma) where the median value reached 16 mm. More detailed inspection suggests that the median grain size increased to c. 11 mm between 4.0 and 3.8 Ma, then increased gradually to 12 mm prior to 2.6 Ma. Median grain size was more or less constant at 12 mm from 2.6 to 2.0 Ma, and again increased rapidly to 16 mm by 1.8 Ma (Fig. 8a). In contrast, the median diameter of the coarser distribution is highly variable around 70 mm, especially before 3.0 Ma and does not show any long-term trend (Fig. 8b). Median diameter of the finer distribution in the conglomerate matrix is variable around 11 mm and does not show a long-term trend until 2.5 Ma, after which it increased slightly (Fig. 8a). Median diameter of the coarser distribution in the conglomerate matrix is highly variable around 90 mm especially before 3.0 Ma and does not show any long-term trend (Fig. 8b). The median diameter of the coarser distribution tends to be larger than that of aeolian siltstones, especially before 3.6 Ma. The finer distribution corresponds to the grain size range characteristic of long-term suspension (,20 mm), whereas the coarser distribution corresponds to the range characteristic of short-term suspension (20–70 mm) and modified saltation (70– 100 mm) (Pye 1987). Thus, we interpret the finer distribution to represent dust transported by longterm suspension, whereas the coarser distribution

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Fine distribution Median diameter (mm) 8

10

12

14

16

(a)

Coarse distribution Median diameter (mm) 20

40

60

80

100 120

(b)

Fine silt distribution Content (%) 0

20

40

60

80

100

(c)

3

Upper Lower

Age (Ma)

2.5

Xiyu Formation

2

Low Mid.

4.5

Upper

4

Artux Formation

3.5

Conglomerate matrix Aeolian siltstone

Wuqia Gr.

8

10

12

14

16

20

40

60

80

100 120

Fig. 8. Variations in: (a) median diameters of the finer distributions; (b) median diameters of the coarser distributions; and (c) contents of the finer distributions in the aeolian siltstone and the matrix of conglomerate of the Artux and Xiyu formations.

covered by aeolian silt, so that the dust source areas were distant and the coarser component could not have been transported by saltation or short-term suspension. 110 R = 0.74

100

Median diameter ( mm)

represents dust transported by short-term suspension and saltation. Because the finer distribution is considered to represent long-term suspension, the upward increase in median diameter of the finer distribution may be interpreted to indicate that the surface wind intensity that lifted the dust from the Taklimakan Desert increased after c. 4 Ma towards 1.8 Ma. The coarser distribution with the median diameter exceeding 80 mm occurs almost exclusively within the lower unit of the Xiyu Formation in which orange to yellow siltstone intercalations are less common thinner and tend to be lenticular. Contents of coarser distribution within the orange to yellow siltstone are close to 100% within this interval (Fig. 8c). We consider that the distribution of aeolian silt on the alluvial fan was patchy and that the coarser component may have been supplied from nearby gobi and wadi. The positive correlation between the content and the median diameter of the coarser distribution also support this idea (Fig. 9). In the upper unit of the Artux Formation, except for the last two cycles, there is generally, a greater quantity of the finer distribution of aeolian siltstone, median diameters tend to be smaller, and sorting tends to be better. This may suggest that the locality was in a distal fan or flood plain setting that was completely

90 80 70 60 50 40 30 0

0.2

0.4

0.6

0.8

1

Content of coarser distribution

Fig. 9. Correlation between the median diameter and the content of coarser distribution of the aeolian siltstones in the Artux and Xiyu formations.

DESERTIFICATION OF THE TARIM BASIN

10

15

20

25

30

35

Upper Lower

3

Upper

2.5

Xiyu Formation

3.5

4

4.5

Lw. Mid.

Isozaki (2009) examined the provenance of orange and yellow siltstone for its fine silt fraction (,16 mm) for selected samples using electron spin resonance (ESR) signal intensity of the E10 centre, an unpaired electron at an oxygen vacancy, of quartz (hereafter we call ESR signal intensity), which is measured after gamma ray irradiation and heating at 300 8C for 15 minutes. This ESR signal intensity represents the number of oxygen vacancies in quartz (Toyoda & Ikeya 1991), which shows a clear positive correlation with the age of the host rock (Toyoda & Hattori 2000). The ESR signal intensity is reset at temperatures of c. 200 8C (Toyoda 1992); this can therefore be used as a thermo-chronometer in an approximate sense. In her study, Isozaki (2009) focused on quartz because quartz is a major component of the aeolian dust and its provenance is likely to be representative of the bulk aeolian dust. Furthermore, quartz is resistant to chemical and physical weathering and diagenesis. She also focused on the fine fraction of ,16 mm because she was especially interested in those dusts that were transported over long distances by long-term suspension. Twenty-six yellow siltstone samples from the Artux and Xiyu formations were analysed for ESR signal intensity. ESR signal intensity was expressed in spin units, where one spin unit is equivalent to 1.3  1015 spins/g (Toyoda & Naruse 2002). Reproducibility of ESR signal intensity measurement was better than +0.9 spin units. Figure 10 shows changes of ESR signal intensity for quartz in the fine silt fraction of orange and yellow siltstone deposited between 4.6 Ma and 1.8 Ma. The ESR signal intensity in the lower and middle units of the Artux Formation (from 4.6 Ma to 4.1 Ma) is between 17.5 and 20.7 (average ¼ 18.7), suggesting a dominant late Precambrian to Palaeozoic age for the source rocks. The ESR signal intensity becomes as high as 30.5 in the upper unit of the Artux Formation and the

5

2

Artux Formation

Provenance of aeolian deposits at Yecheng

ESR signal intensity (spin unit) 0

Age (Ma)

When the median diameter of the finer distribution of the siltstone is compared with that of the conglomerate matrix from the same stratigraphic level, the former is smaller than the latter for sediments deposited before 3.75 Ma, whereas the two are more or less the same after 3.75 Ma. The age of 3.75 Ma roughly corresponds to the time when conglomerate intercalations became more frequent. It is possible that the contribution of recycled dust particles, which were transported to the drainage areas, to the fine fraction of the sediment discharge increased rapidly after around 3.75 Ma.

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Wuqia Gr.

Fig. 10. Changes in ESR signal intensity of quartz in fine silt (,16 mm) fraction of aeolian siltstone in the Artux and Xiyu formations. Also shown are ESR signal intensities of quartz in fine silt fractions of mudstones and sandstones of the underlying Wuqia Group and the late Quaternary mountain loess from Yutian.

basal part of the Xiyu Formation (from 4.0 Ma to 3.5 Ma). Very high ESR signal intensities with the average of 25.3 within this interval suggest a Precambrian age for the source rocks. The ESR signal intensity suddenly decreases to 12.0 at 3.4 Ma and stays at low values between 12.3 and 1.7 with the average of 9.0 thereafter, suggesting Palaeozoic to Mesozoic for source rock ages. The ESR signal intensities of quartz in the fine silt fraction of mudstones and sandstones of the underlying Wuqia Group ranges from 13.0 to 22.7, with an average value of 17.8, which is similar to the values for the lower unit of the Artux Formation. In contrast, ESR signal intensities of quartz in fine silt fraction from ‘mountain loess’ in the Yutian area (Fig. 1) ranges from 5.2 to 7.5 with the average of 6.5 (n ¼ 4), which is similar to values in the upper unit of the Xiyu Formation.

Desertification, dust accumulation and tectonic tilting in the southwestern Tarim Basin 4.6 to 4.0 Ma We show that accumulation of aeolian dust at Yecheng started as early as 4.6 Ma, confirming the

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original suggestion by Zheng et al. (2000, 2003, 2006). The LSR of aeolian siltstone increased gradually from 4.3 Ma and reached its highest values at c. 4.0 Ma. The median grain size of the finer distribution that represents long-term suspended transport is small, being around 9 mm before 4.0 Ma, and suggesting that the intensity of the surface wind was relatively weak and probably not sufficient to elevate the emitted dust to a high enough altitude where it could have been transported by the westerly jet. This is consistent with very low mass accumulation rates of aeolian dust in the North Pacific before 3.6 Ma (Rea et al. 1998). Moderate ESR signal intensity values of 18.7 in the lower and middle units of the Artux Formation are approximately the same as those values of fine silt fraction of mudstones and sandstones in the underlying Wuqia Group (average ¼ 17.8) (Fig. 10), suggesting that the source of the aeolian dusts was probably the same as that in the underlying fluvial sediments, which were delivered by the river flowing westward along the southern margin of the Tarim Basin, based on the palaeocurrent directions at that time (Zheng et al. 2006). The dip of strata is nearly constant for the stratigraphic interval corresponding to this period (Fig. 7a), suggesting no significant tectonic activity at the studied site at that time.

4.0 to 3.6 Ma The LSR of aeolian siltstone increased significantly and became comparable to the present LSRs of ‘mountain loess’ after around 4.0 Ma (Fig. 7b), suggesting development of a desert of significant size in the Tarim Basin. The median grain size of the finer distribution increased from 9 mm at 4.0 Ma to 11 mm at 3.8 Ma (Fig. 8a). This may correspond to the significant increase in aeolian dust mass accumulation rate reconstructed in the North Pacific after c. 3.6 Ma (Rea et al. 1998; Zheng et al. 2004). The orange siltstones that accumulated between 4.0 and 3.8 Ma (in the lower half of the upper unit of the Artux Formation) contain only trace amounts of coarser distribution (Fig. 8c), suggesting that the landscape around the studied site was completely covered by fine silt. A drastic increase in ESR signal intensity seen in the fine silt-sized aeolian dust in the upper unit of the Artux Formation occurred at 4.0 Ma (Fig. 10), which coincides with the onset of northward tilting at Yecheng (Fig. 7a). The increase in ESR signal intensity in the upper unit of the Artux Formation, started approximately at 4.0 Ma (Fig. 10), may reflect switches of the river system around the studied site from westward to eastward flow in response to regional crustal tilting and the consequent changes in the provenance of fluvial detrital

materials. In fact, palaeocurrent directions changed from westward to eastward between the Wuqia Group and the Artux Formation, at c. 4.6 Ma (Zheng et al. 2006), although the timing of the change in direction seems slightly earlier than the onset of increase in ESR signal intensity. Alternatively, the increase in ESR signal intensity may reflect the early stage of unroofing in the mountains surrounding the Tarim Basin where unmetamorphosed sedimentary rocks of Precambrian age or younger sedimentary rocks containing unmetamorphosed detrital quartz grains of Precambrian age were exposed and eroded during the initial stage of surface uplift.

3.6 to 2.6 Ma Slightly lower but still relatively high LSRs of aeolian siltstone continued after 3.6 Ma until 2.6 Ma (Fig. 7b), suggesting the continuous presence of a desert in the Tarim Basin. The median diameter of the finer distribution gradually increased to 12 mm towards 2.6 Ma (Fig. 8a) suggesting gradual intensification of the wind. This is consistent with the observation that mass accumulation rates of aeolian dust increased rapidly from 3.6 to 2.6 Ma in the North Pacific (Rea et al. 1998). Variable but generally low contents of finer sediment in aeolian siltstones during this interval (Fig. 8c) suggest a patchy coverage of aeolian silt on the gobi developed on the alluvial fans, which had probably started to develop at the range front of northern Tibet approximately after 4.0 Ma and reached the studied site at c. 3.6 Ma. Palaeocurrent directions in the middle part of the Xiyu Formation reported by Zheng et al. (2006) also indicate northward flow, supporting the hypothesis of alluvial fan development. A rapid decrease in ESR signal intensity occurred at 3.5 Ma (Fig. 10), which slightly predated the appearance of high-grade metamorphic and plutonic clasts in the Xiyu Formation conglomerate (c. 3.2 Ma). This is reasonable because the resetting of the ESR signal intensity of quartz occurs at c. 200 8C, which is lower than the temperature ranges represented by the high-grade metamorphic rocks and plutonic rocks. Thus, we interpret the dramatic decrease in ESR signal intensity at 3.5 Ma to reflect a change in provenance in the detrital sources. This in turn reflects exposure of low-grade metamorphic strata, in which the quartz ESR signal intensity was reset within the source area (northwestern part of Tibet) at 3.5 Ma as a result of progressive unroofing. ESR signal intensity of quartz gradually decreased from 13 to 10 between 3.6–2.6 Ma (Fig. 10), suggesting a slight increase in the contribution from younger metamorphic quartz caused by continued

DESERTIFICATION OF THE TARIM BASIN 3000

2500

Cgl LSR (m/Ma)

unroofing. This timing approximately corresponds to the time of changes in colour and the geometry of siltstone intercalations from lenses to continuous layers in the lower unit of the Xiyu Formation, at c. 3.4 Ma. This also suggests that the change in colour of the aeolian siltstone reflects a change in its provenance from dominantly unmetamorphosed sedimentary rocks to metamorphic rocks. The tectonic tilting of the study area started from 4.0 Ma and continued until 2.6 Ma, with the peak of the first tilting pulse between 3.2 and 3.0 Ma (Fig. 7a). This timing agrees exactly with the first appearance of high-grade metamorphic and igneous clasts at c. 3.2 Ma. The dominance of plutonic clasts between c. 3.1 and 2.9 Ma reported by Zheng et al. (2006) suggests rapid unroofing and exposure of high-grade metamorphic and plutonic rocks in the drainage area in northwestern Tibet. Sedimentation rates for conglomerate deposition in the Artux and Xiyu formations also peaked between 3.3 and 3.0 Ma (Fig. 7c), consistent with rapid unroofing. Thus uplift and erosion driven by crustal shortening probably started at about 4.0 Ma and peaked at 3.1 Ma at least around Yecheng in northwestern Tibet.

61

2000

1500

1000

500

0 0

50

100

150

200

250

Yel Slt LSR (m/Ma)

Fig. 11. Correlation between LSRs of aeolian siltstone and conglomerate in the Artux and Xiyu formations at Yecheng.

of detrital materials and their supply to the surface of the alluvial fans may be one of the main factors that control the emission of aeolian dust.

2.6 to 2.1 Ma

2.1 to 1.8 Ma

The LSR of aeolian siltstone decreased to a minimum value of 33 m/Ma between 2.6 and 2.1 Ma (Fig. 7b). This decrease in aeolian silt LSR slightly postdates the temporal decrease in aeolian mass accumulation rate (MAR) in the North Pacific Ocean between 2.8 and 2.6 Ma possibly because of poorer age controls in the North Pacific core. This period is also the time when intercalations of yellow siltstone layers are sparse. The median grain size of finer aeolian siltstone remains at approximately the same value of 12 mm during this period (Fig. 8a). A low content of finer fraction in aeolian siltstones (Fig. 8c) suggests a patchy coverage of aeolian silt on the gobi developed on the alluvial fans. The ESR signal intensity of quartz stays around 10 (Fig. 10), indicating continued contribution of younger metamorphic quartz. This interval of low aeolian LSR coincides with an interval characterized by cessation of tectonic tilting and associated decrease in conglomerate LSR (Fig. 7). Because gobi and wadi development on the alluvial fan are considered to be the main sources of aeolian dust, whereas debris flows are the dominant source of detrital materials to the alluvial fans, it is possible that a small supply of detrital material to the alluvial fans would be a cause of low accumulation rates of aeolian silt. In fact, there is a moderate correlation between LSRs of aeolian siltstone and LSRs of conglomerate (Fig. 11). This correlation suggests that production

The LSR of aeolian siltstone increased again after 2.1 Ma towards the top of the Xiyu Formation at c. 1.8 Ma (Fig. 7b). This increase is associated with an increase in median grain size to 16 mm towards the top of the Xiyu Formation (Fig. 8a), suggesting intensification of surface wind strength during this period. In contrast, the MAR of aeolian dust in the North Pacific slightly decreased during this period. The increases in LSR and median grain size of aeolian siltstone correlate with the second pulse of tectonic tilting, which is dated at approximately between 2.2 and 1.5 Ma (Fig. 7a). This event also correlates with increases in dolerite, gabbro, and diorite clasts and decreases in the proportion of volcanic clasts (Zheng et al. 2006) in the conglomerate of the Xiyu Formation, which is dated at 1.9 Ma, as well as with the increase in conglomerate sedimentation rate after 2.1 Ma (Fig. 7c).

The linkage between northern Tibet uplift, alluvial fan development and the sourcing of aeolian dust from the Tarim Basin As discussed above, the depositional history of aeolian siltstone at Yecheng suggests the onset of aeolian dust emission at 4.6 Ma and the production of significant quantities of dust from the Tarim

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Basin starting at c. 4.0 Ma. This timing coincides roughly with the onset of tectonic tilting in the area at 4.0 Ma. Recently Sun and Liu (2006) and Sun et al. (2008) studied an equivalent Neogene sequence at Sanju, c. 110 km to the east of Yecheng, and reported that both the deposition of aeolian siltstone and the tilting of strata started at 5.3 Ma. However, it is difficult to correlate the timing of these two events with those at Yecheng because no detailed lithostratigraphic, magnetostratigraphic, or dip and strike data were collected from the Sanju section. If age estimates at both sections are correct, then the onset of aeolian dust deposition and the tilting of strata are slightly diachronous. In any event, both structures are related to north–south compression ultimately caused by collision of India and Asia. These tectonic activities also seem to be related with the transpressional activity of the Altyn Tagh Fault along which active exhumation occurred around 9–7, 5.15– 4.15 Ma, and 2.11–1.18 Ma according to apatite fission track ages (Zhang et al. 2008). The second period coincides roughly with the onset of desertification of the Tarim Basin between 5.3 and 4.6 Ma (Zheng et al. 2000, 2003, 2006; Sun & Liu 2006; Sun et al. 2008) and the development of a substantial desert by 4.0 Ma. Intensification of surface winds between 4.0 and 3.8 Ma may also be related to this uplift and probably enabled dust to be lifted up to higher altitudes where it can be entrained by the westerly jet and transported long distances into the North Pacific (Zheng et al. 2004). Recent Atmosphere–Ocean coupled General Circulation Model (AOGCM) simulations by Kitoh (2004) demonstrate that desertification of inland Asia at middle latitudes occurs when the height (and extent) of Tibet reaches 60% of the present value, suggesting that there is a threshold for the height (or extent) of Tibet beyond which desertification of Tarim Basin will occur. It is possible that uplift of northern Tibet starting from c. 5 Ma crossed this threshold by c. 4 Ma. The first pulse of tilting and uplift culminated at c. 3.1 Ma, a time associated with enhanced supply of debris flows to alluvial fans and continued high emissions of aeolian dust. The uplift of northern Tibet seems to have weakened between 2.6 and 2.1 Ma, reducing sediment discharge to the alluvial fans that developed on the northern foot of the Kunlun Mountains. The decrease in LSR of alluvial conglomerate at Yecheng between 2.6 and 2.1 Ma may be explained partly by the decrease in East Asian summer monsoon intensity (e.g. An et al. 2001) and the consequent reduction in snow accumulation in Tibet. However, the peak in East Asian summer monsoon intensity is at c. 2.6 Ma, whereas the peak in alluvial conglomerate LSR is

between 3.3 and 3.0 Ma. Thus, the strength of the East Asian summer monsoon does not seem to be the dominant factor in controlling the deposition of alluvial conglomerate adjacent to northern Tibet. In any event, this decrease in supply of detrital material may have caused the reduction of the dust emission from the basin. This study also revealed a second pulse of uplift and tilting from 2.2 to 1.5 Ma. Increased tectonic activity in northern Tibet is also implied by significant flux of pyroclastic rocks near the top of the Xiyu Formation at Yecheng. The LSRs of aeolian siltstone and conglomerate both increased during this period. Our results therefore suggest a possible link between tectonic uplift, erosion and the supply of detrital materials to alluvial fans, followed by enhanced dust emission from gobi developed on the margin of the alluvial fans and from dry lakes developed near the margin of alluvial fans. In addition, the median grain size of the finer distribution increased during this period, suggesting further increase in the surface wind intensity, possibly driven by intensification of the Tibetan atmospheric low during spring to summer. This suggests that the second pulse of plateau uplift further enhanced both dust emission and the surface wind intensity that allowed larger amount of fine, siltsized dust to be emitted from the Tarim Basin and transported downwind.

Summary We conducted a semi-quantitative field examination on the occurrence of aeolian siltstones and alluvial conglomerates in the Neogene fluvial sequence at the Yecheng section on the southwestern margin of the Tarim Basin. Our work confirms the previous observations of Zheng et al. (2000, 2003, 2006) that deposition of aeolian siltstone started at 4.6 Ma. Our study further revealed that the deposition of aeolian siltstone became significant after 4.0 Ma. The high LSR of aeolian siltstone was maintained, until at least 1.8 Ma, but probably until the present, suggesting that the desert had grown to a significant size within the Tarim Basin by 4.0 Ma. Examination of the relationship between the tilting rate of strata, LSR of alluvial conglomerate, and LSR of aeolian siltstone revealed a close link between tectonic uplift, erosion, sediment accumulation on alluvial fans, and dust emission from the surface of alluvial fans on the northwestern margin of Tibet. Two pulses of tectonic activity are identified, dating approximately from 4.0 to 2.6 Ma and from 2.2 to 1.5 Ma. These periods are associated with high LSR of alluvial conglomerate and high LSR of aeolian siltstone. Our result, in addition to recent studies in the surrounding area, combined with AOGCM simulations by Kitoh (2004) further

DESERTIFICATION OF THE TARIM BASIN

supports models that link development of the Taklimakan Desert at c. 5–4 Ma to when the height (and extent) of the Tibetan Plateau exceeded a certain threshold value during the uplift of northwestern Tibet that started around 5 Ma. Our work also suggests continued increase in the surface wind intensity since 4 Ma in association with the progressive uplift of northern Tibet. The timings of desertification at 5–4 Ma and intensification of the surface wind since 4 Ma precede the onset of northern hemisphere glaciation (c. 3 Ma; e.g. Raymo et al. 2006). Consequently, an influence of northern hemisphere glaciations over erosion in this part of central Asia is unlikely. Although further evidence is needed, our results support the idea that the uplift of northern Tibet has caused desertification in the Tarim Basin and the subsequent emission of dust since c. 4 Ma. This research was partly supported by the Mitsubishi Foundation provided to R. Tada, Fujiwara Natural History Foundation provided to Y. Isozaki, and JSPS fellowship and supporting funds provided to Y. Sun. We thank P. Clift and two anonymous reviewers for their careful reading of the manuscript and many constructive suggestions.

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Zheng, H., Powell, C. M., Rea, D. K., Wang, J. & Wang, P. 2004. Late Miocene and mid-Pliocene enhancement of the East Asian monsoon as viewed from the land and sea. Global and Planetary Change, 41, 147–155. Zheng, H., Huang, X. & Butcher, K. 2006. Lithostratigraphy, petrography and facies analysis of the Late Cenozoic sediments in the foreland basin of the West Kunlun. Palaeogeography, Palaeoclimatology, Palaeoecology, 41, 61– 78. Zheng, H., Tada, R., Jia, J. & Lawrence, C. 2010. Cenozoic sediments in the southern Tarim Basin: implications for the uplift of northern Tibet and evolution of the Taklimakan Desert. In: Clift, P. D., Tada, R. & Zheng, H. (eds) Monsoon Evolution and Tectonics – Climate Linkage in Asia. Geological Society, London, Special Publications, 342, 67–78.

Cenozoic sediments in the southern Tarim Basin: implications for the uplift of northern Tibet and evolution of the Taklimakan Desert HONGBO ZHENG1*, RYUJI TADA2, JUNTAO JIA3, COLIN LAWRENCE4 & KE WANG1 1

School of Earth Science and Engineering, Nanjing University, Nanjing, 210093, China

2

Department of Earth and Planetary Science, University of Tokyo, Tokyo 113-0033, Japan 3

State Key Laboratory of Marine Geology, School of Ocean and Earth Science, Tongji University, Shanghai, 200092, China 4

School of Earth and Geographic Sciences, The University of Western Australia, Crawley, WA 6009, Australia *Corresponding author (e-mail: [email protected]) Abstract: Cenozoic sedimentary successions along the southern margin of the Tarim Basin, western China, reach up to 10 km in thickness. The two studied sections, the Yecheng and Aertashi, comprise c. 4.5 km and c. 7.0 km of clastic sedimentary rocks respectively. The base of the Yecheng section has been dated palaeomagnetically to be about 8 Ma. Age control of the Aertashi section is based on 87Sr/86Sr measurements (for the basal marine bed), together with magnetostratigraphy and regional stratigraphic correlation. The lower part of each section is mainly composed of fine-grained mudstone and fine sandstone, which makes up the Wuqian Group (Miocene). The palaeoenvironment is low-energy, meandering and braided streams. The middle part is composed of red mudstone, sandstone with thin conglomerate beds, which make up the Artux Formation (Pliocene). The palaeoenvironment is a distal- to mid-fan environment. The uppermost part of the section, known as the Xiyu Formation (Plio-Pleistocene), consists of cobble and boulder conglomerate intercalated with massive siltstone lenses, which formed as proximal alluvial fan and aeolian deposits. Neogene red beds passing upward into upward-coarsening conglomerate and debris-flow deposits record the change in palaeoslope related to uplift of the northern margin of Tibetan Plateau. The formation of aeolian dunes at c. 8 Ma, and underlying playa lake deposits (as at Aertashi), may indicate an arid, enclosed basin in the southern Tarim after this time. Sedimentological characteristics, together with grain size distribution and geochemistry of siltstone bands in the Xiyu and Artux Formations, point to an aeolian origin. This indicates that the Taklimakan Desert and the regional climate regime may have been fully developed by the Early Pliocene. The onset of aeolian sedimentation in the southern Tarim Basin coincided with uplift of the northern Tibetan Plateau, as inferred from the lithofacies change. Tibetan Plateau uplift resulted in the shift of sedimentary environments northwards into the southern Tarim Basin, and could well have triggered the onset of full aridity in the Taklimakan region as a whole.

Central Asia has experienced dramatic changes in both tectonics and climate during the Cenozoic, including the uplift of Tibetan Plateau, which is one of the most important tectonic events in Earth’s recent history. Climatic changes, as represented by the aridification of the Asian interior at one extreme, and monsoon precipitation at another, are also most prominent. Numerous studies have been carried out in the recent past to document the timing, mode and amplitude of these events, the effects of which operate well beyond the immediate region under consideration. In particular, the causal linkage between Asian tectonics

and climatic changes has long been proposed and explored through either geological records or numerical modelling. For example, Ramstein et al. (1997) carried out numerical modelling on the effects of the retreat of Paratethys to the climate and found that this environmental change alone could have largely increased the aridity of Asian interior and strengthened the Asian monsoon. Many others simulated the impact of Tibetan uplift on the evolution of the Asian monsoon and aridity of Asia, even on the northern hemisphere glaciation (e.g. Zhang et al. 2007). In the last few years, a significant number of investigations of the

From: Clift, P. D., Tada, R. & Zheng, H. (eds) Monsoon Evolution and Tectonics –Climate Linkage in Asia. Geological Society, London, Special Publications, 342, 67– 78. DOI: 10.1144/SP342.6 0305-8719/10/$15.00 # The Geological Society of London 2010.

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uplift of the Himalaya and Tibetan Plateau (HTP) have been carried out. These observations indicate that the HTP uplift proceeded by several steps as a northeastwards progression of the uplift area, and that the Asian monsoon evolution also evolved with several steps as a progressive penetration of the area of enhanced precipitation into the Asian continent, as well as the desertification of Taklimakan and Gobi (e.g. Harrison et al. 1992; Molnar et al. 1993; Matte et al. 1996; Copeland 1997; Zheng et al. 2000; An et al. 2001; Guo et al. 2002; Sun & Wang 2005). Uplift of the northern margin of the Tibetan Plateau resulted in deposition of enormous amounts of sediments into the foreland basins of the southern Tarim. Through a series of fluvial and aeolian processes these sediments were delivered into the basin, and eventually made up the Taklimakan Desert. Subsequently, the desert provides dust, not only for the aeolian deposits along the proximal margins of the basin, but also for the distal Chinese Loess Plateau and remote northern Pacific, thus taking part in major geochemical cycles (Zhang et al. 2003). Previously, we had carried out sedimentological and palaeoenvironmental studies on the Yecheng section of Late Miocene to Pleistocene age (Zheng et al. 2006). Results have shown that the northern margin of the Tibetan Plateau experienced

accelerated uplift during the Early Pliocene, and that widespread aeolian deposition commenced at the same time, indicating the existence or even full development of the Taklimakan Desert. In this study, we expand our record to the latest Eocene/ earliest Oligocene by introducing the Aertashi section, located 100 km to the NW of Yecheng. The aim of this study is to conduct a detailed sedimentological and palaeoenvironmental study on the more complete sedimentary successions in the southern Tarim Basin, and to explore their implications for the tectonic uplift of the northern margin of the Tibetan Plateau and the evolution of the Taklimakan Desert.

Regional geology and stratigraphy The field region is located in the southern Tarim Basin and on the northern margin of the West Kunlun Mountains. The Kunlun Mountains mark the northern margin of the Tibetan Plateau (Fig. 1). The basement of the Tarim Basin is Archaean and Proterozoic crystalline metamorphic rocks of mostly amphibolite to granulite facies. Rock types include gneiss, schist and marble, with minor granite and pegmatite dykes (Zhou 2000). Paleocene to Miocene deposits unconformably overlie Jurassic strata and mainly comprise thick deposits of terrestrial red-bed clastic sedimentary rocks.

Fig. 1. Location map showing the distribution of aeolian facies in the Tarim Basin and studied sections. 1, Aertashi section; 2, Yecheng section; 3, Keliyang section (after Zheng et al. 2003).

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Shallow-marine limestones and calcareous shales were deposited periodically until the Oligocene (Zhou 2000). In a more recent study, Ritts et al. (2008) reported a foraminiferal assemblage typical of an open marine environment in mid-Miocene sediments from the southeastern Tarim Basin, adjacent to the Altun Shan. This implies a much younger age for the final retreat of marine environment in the southern Tarim than previously believed (Ritts et al. 2008). In the Miocene, the whole Tarim Basin subsided and a series of terrestrial red sandstones, shales and mudstones were deposited. Uplift and thrusting of the marginal mountains caused deposition of thick piedmont conglomerates along the basin margins (Zhou 2000). The previously studied Yecheng section is introduced here as a reference (Zheng et al. 2000, 2006). The Yecheng section comprises over 4.5 km of Upper Neogene sedimentary rock (Fig. 2). The base of the section is marked by a bed of fine to medium sandstone, with exposed thickness up to 40 m. Observations of the sandstone bed show very good sorting and cross-bedding. We interpret this sandstone unit as being aeolian in origin, representing a palaeo-sand dune deposit (Zheng et al. 2006). The lower part of the section consists of alternating beds of reddish mudstone, siltstone and fine to medium grained sandstone. Based on regional stratigraphic correlation, the strata are equivalent to the upper part of the Miocene Wuqia Group (Sobel & Dumitru 1997). The middle part of the section contains fine-grained sandstone, orange siltstone, yellowish mudstone and thin layers of fine gravel, which comprise the Artux Formation. The uppermost part of the section is named the Xiyu Formation, and is dominated by pebble to boulder conglomerate intercalated with lenses of pale yellow siltstone. All three units are conformable at this location. The Pleistocene Wusu Formation unconformably overlies the Xiyu Formation to the north of the section. Aeolian loess deposits of late Pleistocene to Holocene age mantle low hills across the whole region (Fig. 1). The principal section of the present study is at Aertashi, which lies along the Yarkand River, about 100 km NW of the Yecheng section (Fig. 1). The locality comprises a complete succession from Upper Cretaceous to Pleistocene. Our study focuses on the Cenozoic sediments, with particular focus on the Neogene sequence (Fig. 3). The total thickness of the Cenozoic sequence is c. 7 km, but the uppermost part (Xiyu Formation) was not mapped in detail. The upper part of Aertashi section correlates with the three formations at Yecheng (namely, the Wuqia Group, the Artux Formation and the Xiyu Formation). The Bashiblake Formation (Yin et al. 2002) marks the base of the studied Aertashi section (Fig. 3).

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Magnetostratigraphy About 2500 palaeomagnetic samples were collected through the Yecheng section with a sampling interval of about 1 m in the Artux Formation and Wuqia Group. For the Xiyu Formation, sampling was only possible from the fine sandstone and siltstone bands (Fig. 2). All samples were thermally demagnetized progressively and their magnetizations were measured using a 2G-cryogenic magnetometer at the University of Western Australia. Previously reported magnetostratigraphy only included the upper 2800 m of the section (Zheng et al. 2000). This study presents the complete palaeomagnetic results, which extend to the very base of the section. The observed magnetostratigraphy can be correlated with Cande & Kent’s (1995) polarity timescale (Fig. 2). The Artux Formation ranges from 4.6–3.5 Ma; the Xiyu Formation has an age between 3.5–1.8 Ma. The c. 1.7 km of red beds exposed in the Wuqia Group is dated palaeomagnetically as being between c. 8.0 and 4.6 Ma. Therefore, the sandstone bed at the base of the section is dated to be slightly older than c. 8 Ma. In the Aertashi section, a total of 1107 samples were collected for magnetic polarity determination from the Wuqia Group, through the Artux Formation, up to the base of the Xiyu Formation. Where possible, sample sites were spaced at 4 m intervals, dependent on the availability of suitable lithologies. The uppermost part of the section (Xiyu Formation) contains substantially fewer beds suitable for palaeomagnetic work. This study only presents the results of the section between 3500 –5500 m, which covers the stratigraphic interval from the upper Wuqia Group and the lower part of the Artux Formation (Fig. 3). The polarity stratigraphy of the Aertashi section is constructed based on its correlation with Cande & Kent’s (1995) polarity timescale (Fig. 3), and with the Yecheng section (refer to Fig. 2). The base of the Artux Formation at Aertashi, defined as the first appearance of gravel beds, is dated to be about 5.8 Ma, which is older than the 4.6 Ma assigned to the same formation at Yecheng. The aeolian sand dunes are slightly older than c. 8 Ma, comparable to the Yecheng section.

Lithofacies and palaeogeographic analysis Facies analysis was carried out on both Yecheng and Aertashi sections to define the sedimentary environments. The objectives of facies analysis are to better define a body of rocks with specific characteristics, based on classification of the sedimentary environments in which the rocks formed. This will further give insight into how tectonics and/or climate

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Fig. 2. Lithostratigraphy and magnetostratigraphy of Yecheng section.

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Fig. 3. Lithostratigraphy and magnetostratigraphy of Aertashi section.

change may have controlled the formation of a sedimentary sequence. Lithofacies was defined following reconnaissance mapping and detailed logging. The whole section was then classified into facies associations based on the lithofacies that were most likely to occur in the succession (Zheng et al. 2006).

Facies associations are groups of facies that occur together and are considered to be environmentally or genetically related. Grouping of several associated facies allows a facies of ambiguous origin to be classified and interpreted, because the combined facies association forms a unique and diagnostic sedimentary environment.

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Yecheng section Six broad facies associations were recognized within the Yecheng section. Detailed logs were made of each sub-facies at a type locality, which were subsequently applied to the more general facies associations. At Yecheng, five facies associations were introduced by Zheng et al. (2006); these include: meandering-stream association, braided-stream association, floodplain association, distal- to mid-fan association, and proximal alluvial-fan association. In this study, we introduce a new facies association for the Yecheng section, which is the aeolian dune association (Figs 2 & 4). This unit marks the base of the profile, with an exposed thickness up to 40 m. The lithofacies is characterized by a thick succession of well-sorted fine sandstones, typically showing steeply dipping (.208) festoon-style cross-beds. The thickness of each cross-bed set is typically between 0.421.5 m. An erosive bounding surface is commonly observed truncating the top of the set. Thin to medium bedded, weakly planarlaminated, well-sorted sandstone is subordinate. The lower part of the Yecheng section (the exposed Wuqia Group) is dominated by the deposition of thick successions of fine-grained terrestrial, fluvial clastics. Sedimentation occurred in meandering and braided streams, and on their flood plains or distal alluvial fan. Floodplains typically slope at 0.138–1.08 (Schumm & Khan 1972) and palaeocurrent indicators show that the palaeoslope graded gently to the north and northwest (Fig. 4). Overall, the depositional environment is one of low energy, with abundant deposition of silt, clay and fine sand. The Artux Formation was deposited in a distalto mid-fan environment. Most of the Artux Formation consists of massive siltstone, fine sand, minor clay and granule beds. The depositional environment is interpreted as having been arid, relative to the underlying Wuqia Group, with little water transport in the area, and significant deposition of silt and fine sand by aeolian processes. The section was close to the uplifted areas, but because of relatively minor uplift and the lack of significant running water to transport much sediment, aeolian and flood events dominated the depositional processes. The uppermost part of the Yecheng section is dominated by pebble to boulder conglomerate typical of alluvial-fan debris-flow deposits. The conglomerate is very poorly sorted, with clasts ranging from granules to boulders .2 m in diameter. There is no evidence of fluvial sorting or reworking of the debris flow conglomerates. Some channelling occurs at the base of individual layers, and reverse grading is common. The Xiyu Conglomerate overlies the Artux Formation conformably

in the studied section, but unconformably in other places (Zhou & Chen 1990).

Aertashi section As mentioned above, the upper part of the Aertashi Section correlates with the Yecheng Section (Fig. 3). The following description therefore focuses on the lithofacies of the lower part of Aertashi Section, which includes three facies associations. These are a shallow marine association, a basin plain (lacustrine) association and a playa lake association. The palaeogeographic evolution of the south margin of the Tarim Basin through time, has been reconstructed based on the lithofacies associations of the two combined sections.

Shallow marine (Early– Late Oligocene) The measured Aertashi section starts from the top of the Bashiblake Formation (Figs 3 & 4). Field examination reveals that the formation is about 200 m thick at this site, and consists of limestone and calcareous shale, with fossil shells up to 20 cm across. Deposition of the Bashiblake Formation along the southern margin of Tarim Basin was widespread. The easternmost exposure occurs at Keliyang, about 100 km to the SE of Yecheng (Fig. 1). Palaeomagnetic measurement of the Aertashi section does not yet go down to the Bashiblake Formation. Instead, 87Sr/86Sr isotope analysis of the shells was carried out to provide an approximate age reference. Three shells yield 87Sr/86Sr values between 0.707744 and 0.707942. When plotted against the Cenozoic global 87Sr/86Sr curve (Palmer & Edmond 1992), these values correspond to ages ranging from 35 –33 Ma. This is in broad agreement with previous estimates, which claimed that the Bashiblake Formation was formed at around 30 Ma (Sobel & Dumitru 1997). Unambiguous evidence for uplift is provided by marine fossils originally deposited at or below sea level, but now lying at elevations of 1700 m. The marine environment formed in a shallow embayment known as the Tadjik Sea, part of the large Turan Sea further to the west (Fig. 5). During the Early Oligocene, the marine basin retreated westward and was replaced by continental environments (Burtman 2000).

Basin plain and fan delta (Early Miocene) The Bashiblake Formation is unconformably overlain by c. 7 km of Neogene sedimentary rocks, although the age of the sediment was relatively poorly constrained (Sobel & Dumitru 1997). Field observation suggests that a lacustrine succession passes up into a fan delta environment (total

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Fig. 4. A schematic diagram showing the composite stratigraphy, palaeocurrents and sedimentary environments of the Cenozoic sediments in the foreland basins of the southern Tarim.

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Fig. 5. Cenozoic evolution of the sedimentary environments of the southern Tarim Basin as inferred from the Yecheng and Aertashi sections.

thickness 300 m). The unconformity between deposition of the fossiliferous marine sediments and development of a fairly deep, maybe 50–60 m, lacustrine basin, must represent a significant time gap. It is proposed that the unconformity observed at the base of the Wuqia Group represents an erosion surface resulting from the regional change in deformation at this time. Apatite fission track ages for Cretaceous and Jurassic sedimentary rocks at Aertashi and Kusilaf indicate partial exhumation at c. 18 Ma (Sobel & Dumitru 1997). The maximum age limit for the base of the Wuqia Group at Aertashi is thus estimated as younger than c. 18 Ma.

Playa lake and distal fluvial (Middle Miocene) This unit is about 300 m thick, and largely contains red-coloured, fine-grained mudstone and siltstone, intercalated with gypsum. The gypsum appears as small crystals, mixed with muddy clastic

sedimentary rocks, or it can be formed as a layer of a few centimetres thick. Laminations within each bed can be observed in places.

Aeolian dunes (Late Miocene, c. 8 Ma) This unit reaches up to 240 m in thickness. The lithofacies is characterized by a thick succession of well-sorted, fine sandstones, typically showing steeply dipping (.208) festoon-style cross-beds. Cross-bedding is typically accentuated by preferential weathering along the bedding planes. The thickness of each cross-bed set is typically between 0.4–1.5 m. An erosive bounding surface is commonly observed truncating the top of the set. Thin to medium bedded, weakly planarlaminated, well-sorted sandstone is subordinate. Two groupings of cross-bed orientations are evident from the inferred aeolian sand unit. The average set thickness is c. 80 cm with festoon style cross-beds typically truncated at the top of the co-set. The mean orientations for the two

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groups are 22/298 (12 measurements) and 30/197 (23 measurements) (Fig. 4). It is assumed that these orientations represent the wings of 3-dimensional aeolian dunes and the bisecting angle c. 2408 represents the unidirectional wind direction forming Barchan style dune. The aeolian strata lie below the magnetostratigraphy analysis performed in this study but correlation indicates that they are older than 7.5 Ma. At the top of the aeolian section, fluvial mudstones interfinger with the aeolian dune sandstones.

Fluvial and distal alluvial (mid- to Late Miocene) The upper Wuqia Group records a change from an aeolian to a flood plain environment. Some of the sandstones have a yellow-green colour resulting from reducing conditions. The colour changes locally along the bed suggesting that it is set after deposition. Sandstone point-counting data indicate a shift towards a higher feldspar content and lower lithic content, so that these rocks are classified as lithic feldspar-rich arenites. Palaeocurrents determined from cross-bed foreset orientations and trough axes suggest that palaeoflow was toward the east to southeast. An upward-coarsening trend is observed, with more active channel deposition at the top of the Wuqia Group.

Fluvial, distal alluvial and aeolian (Early Pliocene) At Aertashi, the Artux Formation is defined by the earliest influx of coarse conglomerate, which signifies the change to a new sedimentary environment. Facies analysis of the Artux Formation indicates a distinct change in the fluvial style from preservation of abundant fine-grained sediment, to rare fines preserved only in overbank or crevasse splay settings. Field mapping also revealed the great variety of subfacies in the Artux Formation. For example, some measured units are characterized by confined sediment gravity flows and shallow braided channels. Conglomerate units typically have a scoured base and may be either clast or matrix-supported. Some are characterized by couplets of conglomerate and sandstone indicating surging debris flow deposition in unconfined areas. The sheet flood deposits and sediment gravity flows indicate an outer alluvial fan environment. The conglomerate is more resistive to erosion and forms prominent dark bands visible in the field. The modal compositions indicate a slightly higher quartz content compared to sandstones in the Wuqia Group. Pebble counts of the conglomerate indicate that most of the clasts are limestone (55%) and 90% of the clasts are

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sedimentary in origin. Limestone clasts in a terrestrial environment tend to break down rapidly and their preservation suggests a proximal source (Mack 1984). Imbrication of clasts indicates palaeoflow was northeasterly. The age limit suggests the base of the Artux Formation is c. 5.8 Ma. The sedimentation rate shows increasingly rapid deposition and averages c. 1200 m Ma21. The upper part of the Artux Formation is characterized by thick mudstones, channelized ephemeral stream and crevasse splay deposition. Conglomerate units gradually fade out as finer grain-size equivalents appear. Mud cracks and ripples are observed and gypsum is common. Mud drapes indicate sedimentation in an ephemeral stream and crevasse splay or overbank deposition. The foredeep basin is probably over-filled at this stage and is perhaps similar to the present day attitude of the Yarkand River (see Fig. 1 for location), where flow is transverse to the range front. The modal composition determined for this part of the section is similar to samples from the lower Artux Formation and is classified as a feldspathic litharenite. Measurable palaeocurrent indicators were rare although asymmetric ripples in some places indicate a flow direction toward the east. Palaeogeographic reconstruction illustrates a return to a flood plain environment as the alluvial fan has retreated to the west, perhaps as sediment supply decreased. Erosion of the fault scarp would have continued, causing an overall reduction in the slope angle. The thickness of the Artux Formation at Aertashi is c. 2000 m compared to just 800 m recorded at Yecheng. The stratigraphic log of Zheng et al. (2000) at Yecheng indicates that the conglomerate within the Artux Formation continues up to the base of the Xiyu Formation. The 1000 m thick flood plain palaeoenvironment seen at the top of the Artux Formation at Aertashi is apparently absent. Pebble counts on conglomerate clasts at Yecheng indicate 50% of clasts are sedimentary in origin, including up to 25% limestone in the Artux Formation. This compares to 90% sedimentary clasts, including 55% limestone in the Artux Formation at Aertashi. Intercalated, pale yellowish grey, very fine grained sandstone and siltstone lenses, within the Xiyu Formation at Yecheng, interpreted to be aeolian in origin, are absent from the section at Aertashi. Bedding dips observed at Yecheng are similar, although the strike is east – west compared to north –south at Aertashi. All these lines of evidence suggest that the depositional environments at the two places were different during the deposition of the Artux Formation. We speculate that the Artux Formation at Aertashi may have been associated with the palaeo-Yarkan River, as the fluvial facies are restricted to this particular site.

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At Yecheng, the Artux Formation was deposited in a distal- to mid-fan environment. A number of thin granule- to cobble-conglomerate beds were deposited by sheet and debris flows, indicating minor, proximal uplift to the south. Most of the Artux Formation consists of massive siltstone, fine sandstone, minor claystone and granule beds. The palaeoenvironment is interpreted as arid relative to that reconstructed for the Miocene Wuqia Group, with little water action in the area, and significant deposition of silt and fine sand by aeolian processes (Zheng et al. 2003). The section was close to the uplifted areas, but because of relatively minor uplift and lack of significant running water to transport much sediment, aeolian and flood events dominated the depositional processes.

Proximal alluvial and aeolian (Late Pliocene – Holocene) During this interval an abrupt change in the palaeoenvironment is recorded in the Xiyu Formation. Over the period since 1.8 Ma, 2.5 km of boulder to cobble conglomerate of the Xiyu Formation was deposited at the Yecheng section. A similar thickness and composition of sediments were laid down at Aertashi. The sediments were derived from the West Kunlun mountains and transported as debris and sheet flows, with only minor stream-channel transport. Deposition occurred on a proximal alluvial fan with slopes of 1 –58 to the north (Summerfield 1991), similar to the slopes observed on modern active alluvial fans in the region (Fig. 5). Deposition was episodic, as debris and sheet floods tend to be infrequent but highly energetic, depositing large thicknesses of sediments in a short period of time. The lack of significant stream action on the alluvial fan allowed silt-sized sediments (in the form of silt bands) to accumulate on the uneven fan-floor during times of low sediment input. The silt is interpreted as loess and similar to what is currently being deposited on the top of the Gobi gravels that rim the Tarim Basin (Zheng et al. 2003).

Aeolian deposition and formation of Taklimakan Desert The formation and evolution of the Taklimakan Desert is a regional response to widespread aridification caused by global climatic changes during the Late Cenozoic. Among other factors (such as global cooling), the uplift of the Tibetan Plateau and induced changes in atmospheric circulation may have provided the major driving forces to this Asian aridification (Zhang et al. 2007). The evolutionary history of the Taklimakan Desert and the

loess deposits in the surrounding regions have great potential for improved understanding of palaeoclimatic changes both on regional and global scales, as well as having implications for the uplift history of the Tibetan Plateau. Zhu (1981) claimed that the aeolian landforms of the Taklimakan Desert were formed during the mid- to late Quaternary. Dong et al. (1991) observed a series of buried aeolian sand units in the Neogene sediments in the Taklimakan Desert and surrounding regions and suggested that the desert may have been in existence since the Pliocene or latest Miocene, that is, much earlier than previously thought. There is little doubt that buried aeolian sand provides direct evidence for the existence of arid landscapes. However, because buried aeolian sand units have been found only in a few scattered sites so far, it is far from conclusive as to whether they represent relics of an aeolian landscape of relatively small scale, or a Late Miocene–Pliocene desert as vast as the Taklimakan is today. Aeolian facies are presently very well developed within and surrounding the Tarim Basin, and three distinct aeolian facies can be recognized from the basin centre to the mountain front (Zheng et al. 2003; Fig. 1). These include: the desert facies which dominates the present day Taklimakan Desert; the Gobi facies which rims the desert; and the sandy loess facies, which is distributed in the hilly piedmont regions of the surrounding mountains. Presently the three facies undergo constant reworking through fluvial and aeolian processes. Sediments shed off the uplifted region are first delivered to the basin by episodic alluvial systems. Fine-grained material is then deflated, resulting in the formation of the Gobi (land surface concentrated with gravels) and desert (sediments concentrated with fine to coarse sand). The deflated fine material is transported, sorted and then re-deposited along the basin margins at relative higher elevations. The loess deposits thus preserve information about the evolution of the desert from which they originally came. Previously we have reported that the silt beds in the Artux and Xiyu Formations at Yecheng were aeolian loess deposits, which suggests that the Taklimakan Desert may have been in existence since the Early Pliocene (Zheng et al. 2003). Sun & Liu (2006) later provided more age constraints on the formation of the desert, which suggest that the desert has been in existence since 5.3 Ma. Our new stratigraphic data at both Yecheng and Aertashi indicate that aeolian dunes first developed at about 8 Ma. The aeolian dune unit at Aertashi occurred after a thick succession of playa lake deposits (thick accumulation of gypsum), representing hypersaline conditions. We believe that the combination of playa lake deposition and aeolian

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dunes indicate that the area was very arid at that time, although this line of evidence is insufficient to conclude whether or not this represents a landscape like the Taklimakan today. Strong evidence of severe deflation and aeolian deposition during Plio-Pleistocene lies in the Artux and Xiyu Formations, which contain abundant silt bands of aeolian origin. As discussed above, the Artux Formation at Yecheng is dominated by pale yellow medium- to fine-grained sandstone units and orange siltstone beds, with minor layers of fine gravel. The Xiyu Formation at both sections is composed of pebble to boulder conglomerate intercalated with bands or lenses of pale yellow siltstone. Sedimentological examination, together with grain size analysis and geochemical investigations all prove that these siltstone bands are aeolian in origin, just like the Pleistocene –Holocene loess that is widely distributed in the region (Zheng et al. 2003). The PlioPleistocene loess was deposited episodically on the uneven fan-floor before being buried by the next debris or sheet flow. Widespread loess deposition along the margins of the Tarim Basin indicates that the Taklimakan Desert must have covered a large area during deposition of the Xiyu Formation, and that the climatic regime may have also been similar to that seen today. The onset of aeolian sedimentation in the southern Tarim Basin at c. 8 Ma coincided with the lithofacies change from fine-grained mudstone and sandstone to coarse clasts. We interpret the onset of fully arid conditions in the Tarim region to be the result of the re-organization of the general atmospheric circulation induced by uplift of the northern part of Tibetan Plateau. The forcing of the Late Cenozoic climate by plateau uplift is yet to be fully explored. General Circulation Model (GCM) sensitivity tests run by Ruddiman & Kutzbach (1989) have suggested that the Tibetan Plateau indeed played an important role in inducing the ‘late Cenozoic climatic deterioration’ in the Northern Hemisphere that culminated in the PlioPleistocene ice age.

Conclusions The studied sections at Yecheng and Aertashi, are situated at the northern foot of the West Kunlun Mountains, and comprise 4.5 km and 7.0 km of Cenozoic sedimentary strata, typical of the southern Tarim Basin. The Cenozoic successions preserve nearly continuous records of the sedimentary response to uplift of the northern margin of Tibetan Plateau and palaeoenvironmental evolution in the Tarim Basin. The Aertashi section comprises c. 7 km of non-marine Palaeogene – Neogene strata unconformably overlying Oligocene marine strata of the

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Bashiblake Formation. The c. 2200 m thick Wuqia Group and c. 2000 m thick Artux Formation are conformably overlain by .2500 m of grey, cobble to boulder, polymictic conglomerate of the Xiyu Formation. Facies analysis of the section reveals a palaeogeographic evolution starting from shallow marine conditions in the Early Oligocene, passing to basin plain and fan delta during the Early Miocene. A playa lake environment was developing during the Middle Miocene, and may have culminated in an aeolian dune landscape around 8 Ma, as observed at both Aertashi and Yecheng. The Yecheng Section correlates with the upper part of the Aertashi Section, except that the Artux Formation at Aertashi was formed in an environment with significant channels and running water, which may indicate the presence of a palaeo-Yarkan River, whereas the Artux Formation at Yecheng was deposited in a distal- to mid-fan environment. Deposition of the thick, pebble to boulder conglomerate of the Xiyu Formation marked a phase of strong mountain uplift, although climate change may also have contributed to the environmental change. Development of aeolian dunes at c. 8 Ma, and the hypersaline deposition below, indicates that the regional climate regime was then significantly arid. Deposition of aeolian silt in the Artux and Xifu Formations indicates that the Taklimakan Desert may have been fully developed by that time. The climate regime had evolved so that the desert was able to provide aeolian dust that was deposited episodically on the uneven surface of the alluvial fan before being buried by the next debris or sheet flow. This work was financially supported by a NSFC grant (40830107) and Mitsubishi Foundation awarded to RT.

References An, Z. S., Kutzbach, J. E., Prell, W. L. & Porter, S. C. 2001. Evolution of Asian monsoons and phased uplift of the Himalaya-Tibetan plateau since late Miocene times. Nature, 411, 62– 66. Burtman, V. S. 2000. Cenozoic crustal shortening between the Pamir and Tien Shan and a reconstruction of the Pamir– Tien Shan transition zone for the Cretaceous and Palaeogene. Tectonophysics, 319, 69–92. Cande, S. C. & Kent, D. V. 1995. Revised calibration of the geomagnetic polarity timescale for the Late Cretaceous and Cenozoic. Journal of Geophysical Research, 100, 6093– 6095. Copeland, P. 1997. The when and where of the growth of the Himalaya and the Tibetan Plateau. In: Ruddiman, W. F. (ed.) Tectonic Uplift and Climate Change. Plenum Press, New York, 19–40. Dong, G., Chen, H., Jing, J. & Wang, Y. 1991. Cenozoic paleo-eolian sands in the south marginal area of the Taklimakan Desert. In: Dong, G., Chen, H., Jing, J. &

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Wang, Y. (eds) Glaciers and Environments of West China. China Association of Quaternary Research, Science Press, Beijing, 42–48. Guo, Z. T., Ruddiman, W. et al. 2002. Onset of Asian desertification by 22 Myr ago inferred from loess deposits in China. Nature, 416, 159– 163. Harrison, T. M., Copeland, P., Kidd, W. S. F. & Yin, A. 1992. Rising Tibet. Science, 255, 1663– 1670. Matte, P., Tapponnier, P. et al. 1996. Tectonics of Western Tibet, between the Tarim and Indus. Earth and Planetary Science Letters, 142, 311 –330. Molnar, P., England, P. & Martinod, J. 1993. Mantle dynamics, uplift of the Tibetan Plateau, and the Indian monsoon. Reviews of Geophysics, 31, 3570– 396. Palmer, M. R. & Edmond, J. M. 1992. Controls over the strontium isotope composition of river water. Geochimica et Cosmochimica Acta, 56, 2099– 2111. Ramstein, G., Fluteau, F., Besse, J. & Joussaume, S. 1997. Effect of orogeny, plate motion and land-sea distribution on Eurasian climate change over the past 30 million years. Nature, 386, 788 –795. Ritts, B. D., Yue, Y., Graham, S. A., Sobel, E. R., Abbink, O. A. & Stockli, D. 2008. From sea level to high elevation in 15 million years: uplift history of the northern Tibetan Plateau margin in the Altun Shan. American Journal of Science, 308, 657 –678. Ruddiman, W. F. & Kutzbach, J. E. 1989. Forcing of late Cenozoic Northern Hemisphere climate by plateau uplift in Southern Asia and the American West. Journal of Geophysical Research, 94, 18 409–18 427. Schumm, S. A. & Khan, H. R. 1972. Experimental study of channel patterns. Geological Society of America Bulletin, 83, 1755– 1770. Sobel, E. R. & Dumitru, T. A. 1997. Thrusting and exhumation around the margins of the western Tarim basin during the India-Asia collision. Journal of Geophysical Research, 102, 5043– 5063. Summerfield, M. A. 1991. Global Geomorphology: An Introduction to the Study of Landforms. New York, Wiley.

Sun, J. M. & Liu, T. S. 2006. The age of the Taklimakan Desert. Science, 312, 1621. Sun, X. J. & Wang, P. X. 2005. How old is the Asian monsoon system? Palaeobotanical records from China. Palaeogeography, Palaeoclimatology, Palaeoecology, 222, 181 –222. Yin, A., Rumelhart, P. E. et al. 2002. Tectonic history of the Altyn Tagh fault in northern Tibet inferred from Cenozoic sedimentation. Geological Society of America Bulletin, 114, 1257– 1295. Zhang, X. Y., Gong, S. L., Zhao, T. L., Arimoto, R., Wang, Y. Q. & Zhou, Z. J. 2003. Sources of Asian dust and role of climate change versus desertification in Asian dust emission. Geophysical Research Letters, 30, 2272, doi: 10.1029/2003 GL018206. Zhang, Z., Wang, H., Guo, Z. & Jiang, D. 2007. What triggers the transition of palaeoenvironmental patterns in China, the Tibetan Plateau uplift or the Paratethys Sea retreat? Palaeogeography, Palaeoclimatology, Palaeoecology, 245, 317–331. Zheng, H., Powell, C., An, Z., Zhou, J. & Dong, G. 2000. Pliocene uplift of the northern Tibetan Plateau. Geology, 28, 715 –718. Zheng, H., Powell, C., Butcher, K. & Cao, J. 2003. Late Neogene aeolian loess deposition in southern Tarim Basin and its palaeoenvironmental significance. Tectonophysics, 375, 49–59. Zheng, H., Huang, X. & Butcher, K. 2006. Lithostratigraphy, petrography and facies analysis of the Late Cenozoic sediments in the foreland basin of the West Kunlun. Paleogeography, Paleoclimatology, Paleoecology, 241, 61– 78. Zhou, Z. Y. (ed.) 2000. Stratigraphy of the Tarim Basin. Science Press, Beijing. Zhou, Z. Y. & Chen, P. J. (eds) 1990. Biostratigraphy and Geological Evolution of Tarim Basin. Science Press, Beijing. Zhu, Z. 1981. The Aeolian Landforms of the Taklimakan Desert. Science Press, Beijing.

Non-stationary response of Plio-Pleistocene East Asian winter monsoon variation to ice volume forcing YOUBIN SUN1*, ZHISHENG AN1 & STEVEN C. CLEMENS2 1

State Key Laboratory of Loess and Quaternary Geology, Institute of Earth Environment, Chinese Academic of Sciences, Xian 710075, China

2

Department of Geology, Brown University, Providence, RI 02912-1846, USA *Corresponding author (e-mail: [email protected])

Abstract: Recent progress both in studies of Chinese loess and deep-sea sediments have provided robust and longer records of winter monsoon variation and ice volume change back to the Late Miocene. However, when and how the winter monsoon became coupled with global ice volume change remains uncertain. Here we compare quartz grain size (a reliable winter monsoon proxy) generated from two loess-palaeosol and red clay sequences with a stacked benthic d18O record (a global ice volume proxy). Our results indicate that at longer (.500 ka) timescales, the winter monsoon became strongly coupled with global ice volume change at 2.1 Ma, while at orbital timescales the winter monsoon variations started to be influenced by global ice volume change at c. 3.3 Ma. Correlation coefficients between these two records further indicate that winter monsoon intensity was strongly coupled with global ice volume change during intervals of 0– 1.1 Ma, 1.4– 1.8 Ma, and 2– 2.85 Ma. In spite of these close connections amplitude mismatches between these two records are evident in terms of both long-term trend and glacial–interglacial fluctuations, suggesting that additional processes might have played a role in modulating the response of the winter monsoon variation to ice volume forcing.

East Asian monsoon evolution is dynamically linked to changes in solar insolation and internal boundary conditions (e.g. Prell & Kutzbach 1987; An 2000; Wang et al. 2008), and is closely affected by the integrated behaviours of the global climatic system (i.e. interactions between the global atmosphere, ocean, land and ice systems) (Webster et al. 1998; Wang 2006). In addition, both numerical climate modelling studies and geological evidence suggest that uplift of the Tibetan Plateau has played an important role in enhancing monsoon variability (Prell & Kutzbach 1997; An et al. 2001; Liu & Yin 2002; Clift et al. 2008), possibly resulting in global cooling and eventual development of major Northern Hemisphere glaciation (e.g. Kutzbach et al. 1993; Ruddiman 1997). The history and variability of the East Asian palaeomonsoon has been studied extensively using widespread aeolian deposits in north China (e.g. Liu 1985; Liu & Ding 1998; An 2000). Variations of magnetic susceptibility generated from Chinese loess suggest that changes in the intensity of the summer monsoon show both similarities in climatic cycles and differences in amplitudes of climatic fluctuations relative to global ice volume (Kukla & An 1989; An et al. 1990; Bloemendal et al. 1995). By contrast, variation of the winter

monsoon as revealed by loess grain size indicates stepwise coupling to global ice volume change (Liu & Ding 1993; Ding et al. 1995). At orbital timescales, magnetic susceptibility and grain size of Chinese loess exhibit a dominant 100-ka cycle over the last 0.8 Ma (Ding et al. 1995; Liu et al. 1999; Lu et al. 2004; Sun et al. 2006a) consistent with the periodicity revealed by the deep-sea d18O record (Imbrie et al. 1984; Lisiecki & Raymo 2005). By contrast, absolutedated oxygen isotope records (a proxy for monsoon precipitation) from cave stalagmites reveal a dominant 23 ka cycle over the past 224 ka that is almost synchronous with summer insolation at 658N (Wang et al. 2008). Such a difference in the periodicity suggests that the forcing mechanisms of the summer and winter monsoons are not necessarily similar. Most likely, the winter monsoon (northwesterly wind) is tightly coupled with the high-latitude climate (Ding et al. 1995), whereas the summer monsoon (rainfall) is strongly related to the low-latitude climate as well as crossequatorial moisture transport from the Southern Hemisphere (An 2000; Wang et al. 2008). To better assess the forcing mechanisms of the East Asian monsoon, it is critical to differentiate the impacts of regional tectonic events and global

From: Clift, P. D., Tada, R. & Zheng, H. (eds) Monsoon Evolution and Tectonics –Climate Linkage in Asia. Geological Society, London, Special Publications, 342, 79– 86. DOI: 10.1144/SP342.7 0305-8719/10/$15.00 # The Geological Society of London 2010.

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climate change on the long-term trend and glacial-interglacial variability of the winter and summer monsoons. Comparison of Milankovitch periods between continental loess and deep-sea records indicates that at about 1.6 Ma the ice sheets in the Northern Hemisphere may have reached a critical size, sufficient to modulate changes in the global climate system (Liu et al. 1999). Further comparison of a stacked loess grain size record with a composite marine d18O record shows that for the past 1.8 Ma, the loess-palaeosol record can be correlated almost cycle-to-cycle with the marine record (Ding et al. 2002). However, comparisons of quartz grain size data and the Zr/Rb ratio with deep-sea d18O record reveal that good land-ocean correlation can be extended back to 2.6 Ma (Sun & An 2004; Liu et al. 2004). Since previous studies were based mainly upon visual inspection and didn’t separately evaluate the impacts of global ice volume on the long-term trend and glacial-interglacial variability of the winter monsoon, when (1.6 Ma or earlier?) and how (stepwise or gradually?) the winter monsoon became coupled with global ice volume change remains unclear. Here we present the results of a comparative time-domain analysis of proxies of the East Asian winter monsoon and global ice volume for the past 3.6 Ma. Unlike previous studies, we employ straightforward statistical analyses (i.e. filtering and correlation analysis) of the data sets to define intervals of strengthened and weakened relationship between the two signals on tectonic and orbital timescales. The objectives of this study are twofold: the first is to discern similarities and differences between these two records, particularly in terms of the long-term trend and glacial-interglacial variability, and the second is to evaluate when and how the East Asian winter monsoon became coupled with global ice volume variability. Finally, we tentatively address the relative roles of solar radiation, global ice volume and regional tectonic events in driving the long-term strengthening trend and increasing glacial-interglacial amplitude of the Plio-Pleistocene winter monsoon variation.

Data description and processing Chinese loess is a dust deposit transported by northwesterly winter monsoon winds (Liu 1985; An et al. 1991). The bulk grain size of Chinese loess deposits provides an approximate measure of winter monsoon wind strength; larger bulk grain size indicates the stronger winter monsoon winds (An et al. 1991; Xiao et al. 1992). However, pedogenesis, a post-depositional weathering process, can modify bulk grain size composition of loess-palaeosol and red-clay samples to various degrees (Porter & An

1995; Xiao et al. 1995; Sun et al. 2000). Recent studies suggest that the grain size of quartz particles is a more reliable indicator of winter monsoon strength than is bulk grain size particularly for Late Cenozoic Red Clay deposits (Sun et al. 2006a, b). Here the mean grain size of quartz of two loess-palaeosol and red clay sequences from the central Chinese Loess Plateau, rather than the bulk grain size record of loess-palaeosol sequences, is employed to evaluate on the linkage between the winter monsoon variation and global ice volume forcing. Descriptions of the chronology and palaeoclimatic implications of the two aeolian sequences were detailed in Sun et al. (2006a), and the quartz grain size data are available at the World Date Center for palaeoclimatology (http://www.ngdc. noaa.gov/paleo/paleo.html). Foraminiferal d18O is a function of the temperature and d18O of the water in which it forms (Craig 1957; Shackleton 1967). Benthic d18O records should produce a better ice volume signal than planktonic records because the deep ocean is more uniform in temperature and salinity relative to surface water. Owing to the observed similarity of most marine d18O records and the global nature of the ice volume signal, benthic d18O record also serves as a primary proxy for global ice volume change (Imbrie et al. 1984; Martinson et al. 1987), even prior to the Pleistocene (Shackleton et al. 1995; Zachos et al. 2001). Recently, benthic d18O records from 57 globally distributed sites were compiled to generate a stacked benthic d18O curve (Lisiecki & Raymo 2005). Stacking strategies as well as uncertainties in the stacked curve were described in Lisiecki & Raymo (2005). While the stack certainly incorporates global deep ocean temperature change, the benthic d18O signal remains the best proxy for Plio-Pleistocene global ice volume change. Comparison of quartz grain size with stacked d18O reveals notable mismatches around 1.25 Ma and 2.6 Ma (Fig. 1). Transient coarsening of the quartz grain size indicates that during these two intervals, either the winter monsoon intensity was exceptionally strong or aridity and the geographic range of dust source areas were significantly expanded (Xiao & An 1999; Ding et al. 2002, 2005). Since global ice volume did not exhibit corresponding increases at these two time intervals, the marked coarsening of quartz grain size is probably related to changes in regional climatic/tectonic conditions, for example, stepwise uplift in the northeastern margin of the Tibet Plateau (Sun & Liu 2000; An et al. 2001). To compare the quartz grain size with the stacked d18O record, we truncated the quartz grain size values around these two short intervals. Since chronology of these two records was generated by orbital tuning, the error

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Fig. 1. Variations of quartz grain size (Sun et al. 2006a) and stacked benthic d18O (Lisiecki & Raymo 2005) during the past 3.6 Ma. Shaded bars denote remarkable coarsening of the quartz grain size in L15 and L33.

was estimated to be less than 10 ka (Lisiecki & Raymo 2005; Sun et al. 2006a). For the convenience of comparison, these two data sets were then normalized and interpolated linearly at 2 ka intervals. Both the quartz grain size and stacked d18O records exhibit distinct orbital-scale variations superimposed on long-term trends. Since the longterm trend and glacial-interglacial variation might be related to different forcing mechanisms, it is necessary to separate these two components from the original data sets. Previous studies reveal that loess grain size and deep-sea d18O records are characterized by cyclic fluctuations at earth-orbital eccentricity (400 and 100 ka), obliquity (41 ka), and precession bands (23 and 19 ka) (Liu et al. 1999; Lu et al. 2004; Lisiecki & Raymo 2005; Sun et al. 2006a). Thus, a low-pass filter (.500 ka) is employed to separate the long-term trend from these two data sets, and the residual can be regarded as an expression of the orbital-induced variability.

Comparison of quartz grain size and benthic d18O record Visual comparison of quartz grain size with benthic d18O record suggests that although the winter monsoon and global ice volume exhibit similar variability at glacial-interglacial timescales, the degree of coupling varies over the last 3.6 Ma (Fig. 2). To assess coupling between quartz grain size and benthic d18O, we calculated the timevarying correlation coefficient between quartz grain size and stacked d18O using a 200 ka window with 20 ka time steps (e.g. 0– 200 ka, 20 – 220 ka, etc.). Correlation coefficients indicate that the East Asian winter monsoon strength is variably coupled with global ice volume change: (1) prior

to 3.3 Ma, the winter monsoon was decoupled with global ice volume change; (2) during intervals of 2.85– 3.3 Ma, 1.8–2 Ma, and 1.1–1.4 Ma, winter monsoon was weakly coupled with global ice volume change; (3) between 2– 2.85 Ma, 1.4– 1.8 Ma, and 0–1.1 Ma, the winter monsoon was strongly coupled with global ice volume change (Fig. 2). Prior to 3.3 Ma, the winter monsoon intensity is characterized by small-amplitude fluctuations and an abrupt weakening around 3.4 Ma, whereas the d18O record exhibits a gradual increasing trend with glacial-interglacial variations superimposed (Fig. 2). Cycle-by-cycle correlation between these two records is less clear and correlation coefficients are less than 0.1 within this interval. Between 2.85 and 3.3 Ma, correlation coefficients between these two records range from 0.2 to 0.4, indicating that winter monsoon intensity started to become weakly coupled with global ice volume change at glacial-interglacial timescales. The quartz grain size exhibits distinct glacial-interglacial variations with rather uniform amplitude, whereas the d18O record displays similar glacial-interglacial variations, superimposed on a gradually increasing trend. Strong winter monsoon circulations as indicated by five quartz grain size troughs roughly correspond to ice maxima. After 2.85 Ma, the quartz grain size exhibits large-amplitude variations particularly after 1.2 Ma, corresponding well with glacial-interglacial fluctuations of the d18O record (Fig. 2). Correlation coefficients (.0.4 for most intervals) indicate that the winter monsoon intensity is strongly related to global ice volume change, with the exceptions of two short intervals (1.8–2 Ma and 1.1–1.4 Ma). During 1.8–2 Ma, quartz grain size indicates that the winter monsoon intensity is relatively weak compared to previous glacial-interglacial cycles,

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Fig. 2. Comparison of normalized quartz grain size with stacked benthic d18O and their time-varying correlation. High quartz grain size indicates a weak winter monsoon during interglacial times, whereas high d18O record denotes decreased ice volume during warm periods. Dashed lines indicate the cycle-to-cycle correlation of strong winter monsoons and ice maxima.

whereas the d18O record suggests that the global ice volume continues to increase gradually. The reason for this decreased coupling is not well understood, possibly due to weakening of ice rafting events in the Nordic Seas (Jansen et al. 2000). Between 1.1–1.4 Ma, weak coupling between the winter monsoon and global ice volume is likely attributed to the remarkable coarsening of the quartz grain size around 1.25 Ma. Notably, correlation coefficients between these two records gradually

increased during 0.6–1.1 Ma, and subsequently reached a maximum and became stable for the last 0.6 Ma. Comparison of the long-term trends of these two records can provide a reasonable evaluation of the relationship between the winter monsoon and global ice volume change at tectonic timescales. The long-term trend of global ice volume increased gradually over the interval of 0.9–3.6 Ma and became stable over the last 0.9 Ma (Fig. 3). In

Fig. 3. Long-term trends of the winter monsoon variation (arrow lines) and global ice volume change after removing high-frequency components (,500 ka) of the quartz grain size and d18O records.

WINTER MONSOON VARIATION AND ICE VOLUME FORCING

contrast, the winter monsoon evolution is characterized by slight weakening during 3.1–3.6 Ma and then sharp strengthening around 2.5–2.9 Ma. Between 2.1–2.4 Ma, the winter monsoon intensity was weakened again. Subsequently, the winter monsoon intensity strengthened gradually until 0.9 Ma. After 0.9 Ma, the long-term trend in winter monsoon intensity stabilized. It is obvious that the long-term trends of these two records are significantly different before 2.1 Ma and match very well thereafter.

Forcing mechanisms of the Plio-Pleistocene winter monsoon variation Mechanisms driving the Plio-Pleistocene winter monsoon variation include external forcing and internal forcing (e.g. changing boundary conditions and interactions among different climate components) (e.g. An et al. 1990; Liu & Ding 1993; Ding et al. 1995; Guo et al. 1996). At tectonic timescales, global climate change and regional tectonic events might have had different impacts on the East Asian monsoon variation (Xiao & An 1999; An 2000; Sun & Liu 2000). For example, increase in the global ice volume will result in strengthening of the winter monsoon and weakening of the

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summer monsoon, whereas tectonic uplift (e.g. the Qinghai-Tibet Plateau) would lead to simultaneously strengthening of the winter and summer monsoons (An et al. 2001). More recently, comparison of loess grain size and deep-sea records suggests that strengthening monsoonal circulation since the Late Miocene was associated with progressive development of global ice sheets, while abrupt changes occurred at 7.6, 3.4, 1.2–0.9 and 0.6 Ma are interpreted as the consequence of the episodic uplift of the Tibetan Plateau (Sun et al. 2008). Our results suggest that, with regard to the long-term trend, the winter monsoon evolution was decoupled with global ice volume before 2.1 Ma, and subsequently the winter monsoon was strongly coupled with the ice volume change. Shift of the coupling relationship around 2.1 Ma might be attributable to the dynamic links between tectonic uplift and climate change, since tectonic changes (e.g. enhanced uplift of the northeastern Tibetan Plateau) may have brought ice volume to a critical threshold, at which point expansion in the ice volume triggered more extensive winter monsoon fluctuations (Liu et al. 1999). After removing the long-term trend, differences in the glacial-interglacial amplitude of these two records are clearly expressed (Fig. 4). As evidenced by the detrended quartz grain size variation, the

Fig. 4. Comparison of: (a) 658N July solar insolation (Berger & Loutre 1991); (b) detrended quartz grain size; (c) detrended d18O; and (d) the amplitude difference between the detrended quartz grain size and d18O records. Solid and dashed lines denote the approximately average amplitude of winter monsoon and global ice volume during interglacial and glacial periods, respectively.

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winter monsoon intensity during interglacial times was relatively stable over the interval 1.25– 3.6 Ma with stepwise weakening around 1.25 and 0.5 Ma. During glacial times, however, the winter monsoon intensity increased in two stepwise increments around 2.65 and 0.65 Ma. Similar to the quartz grain size variation, detrended d18O indicates that during interglacial times, global ice volume was decreased sharply around 1.25 and 0.42 Ma, whereas during glacial periods global ice volume exhibits two stepwise increases around 2.65 and 0.65 Ma. Besides the above-mentioned similarities, discrepancies between these two records are also notable. First, several exceptional quartz grain size spikes (L9, L15, L32, L33 in Fig. 4b) do not match the rather uniform amplitude of the ice volume during the corresponding glacial times. Second, a notable decrease in the interglacial winter monsoon intensity occurred around 0.5 Ma, significantly earlier than the beginning of the largestamplitude d18O fluctuations around 0.43 Ma. These discrepancies suggest that the response of winter monsoon variation to ice volume forcing might be disturbed occasionally by additional processes, for example, regional tectonic events (Sun & Liu 2000; Zheng et al. 2004), desert expansion (Ding et al. 2005), and even Southern Hemisphere climate (Liu et al. 1996; Guo et al. 2009). The amplitude difference, estimated by quartz grain size minus d18O (Fig. 4), likely reflects the forcing-response relationship between global ice volume and the winter monsoon. At tectonic timescales, the amplitude difference can be divided into three intervals: (1) relatively small between 2.65 and 3.6 Ma, (2) moderate during 2.65– 1.25 Ma, and (3) significant after 1.25 Ma. Variation in the Northern Hemisphere summer solar radiation received at high-latitude, which has been considered as the primary factor driving the glacial-interglacial climate fluctuations (Hays et al. 1976; Kutzbach 1981), didn’t exhibit significant change around 2.65 and 1.25 Ma. Thus, the large shifts in winter monsoon intensity are not due to the external/ orbital forcing, but rather to internal forcing. Taking into account above mentioned coupling relationship, we infer that the first shift (2.65 Ma) was induced by phased uplift of the northeastern Himalaya-Tibetan Plateau and the onset of the Northern Hemisphere glaciation (e.g. An et al. 2001; Clemens et al. 2008), whereas the second shift (1.25 Ma) was related to the full development of the Northern Hemisphere ice sheet and to the commencement of the mid-Pleistocene transition (Lisiecki & Raymo 2007). At orbital timescales, variation of the amplitude difference reveals that during the deglaciations (i.e. glacial/interglacial transitions), the winter monsoon intensity weakened slowly while the global ice

volume decreased abruptly. Negative amplitude difference suggests that the winter monsoon intensity is stronger than expected during deglaciations. During the interglacial/glacial transitions, global ice volume increased sharply in contrast to a gradual strengthening of the winter monsoon intensity. Positive amplitude difference indicates a prolonged weakening of the winter monsoon during interglacial/glacial transitions and even early glacial times, though global ice volume has already increased. The significant amplitude difference between these two records suggests that during the climatic transitions the response of the winter monsoon to ice volume forcing is nonlinear and might be modulated by complicated feedbacks within the earth climate systems (e.g. ice albedo, vegetation cover, greenhouse gas concentrations) (King 1996; Shackleton 2000).

Summary Comparison of quartz grain size with stacked d18O, together with time-varying correlation analysis, indicates that coupling of the winter monsoon variation to ice volume forcing was non-stationary over the last 3.6 Ma. At longer (.500 ka) timescales, the winter monsoon became strongly coupled with global ice volume change after 2.1 Ma, whereas before 2.1 Ma the long-term trend of winter monsoon intensity was decoupled with ice volume change, probably disturbed by regional tectonic/ environmental events. At orbital timescales, the East Asian winter monsoon variation is coupled with global ice volume change in a non-stationary fashion: decoupled prior to 3.3 Ma, weakly coupled over the intervals of 1.1–1.4, 1.8–2 and 2.85 –3.3 Ma, and strongly coupled during other time intervals (0–1.1, 1.4–1.8 and 2– 2.85 Ma). Amplitude mismatches between these two records are evident in terms of both long-term trend and glacial-interglacial fluctuations, implying that additional processes (e.g. regional tectonic/ environment events, feedbacks within the climate systems) might be involved in modulating the response of the winter monsoon variation to ice volume forcing. Two anonymous reviewers gave insightful comments on this manuscript. This work was supported by the National Basic Research Program of China (No. 2010CB8334003) and the ‘One-hundred Talents’ programme of the Chinese Academy of Sciences.

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Radiometric dating of the late Quaternary summer monsoon on the Loess Plateau, China T. STEVENS1* & H. LU2 1

Department of Geography, Royal Holloway, University of London, Egham, Surrey, TW20 0EX, UK 2

School of Geographical and Oceanographical Sciences, Nanjing University, Nanjing 210093, China *Corresponding author (e-mail: [email protected])

Abstract: Recent advances in radiometric dating have enabled independent investigation into monsoon variations. In this study, summer monsoon pedogenesis proxies (CaCO3 and magnetic susceptibility) have been analysed by optically stimulated luminescence (OSL) age for five loess-palaeosol sections over the Chinese Loess Plateau. The use of CaCO3 is complicated by the multiple influences on its variation. However, changes in magnetic susceptibility can be used as a proxy for summer monsoon induced pedogenesis. The data suggest that the summer monsoon in north-central China is not prone to high frequency shifts, although abrupt transitions occur. The overall patterns show general decreasing trends from c. 50 to 18 ka. However, between 9 and 6 ka, magnetic susceptibility increases abruptly and dramatically at the sites. These findings suggest that the Holocene ‘optimum’ in the region may be a more recent phenomenon than previously suggested, and that this summer monsoon intensity increase significantly post-dates the insolation peak occurring at 11.5 ka. An apparent close correspondence to ice volume is suggested to be a consequence of forcing via atmospheric circulation. Independently dated records that employ high sampling resolution can be used to test this hypothesis, together with suggestions over the apparent lag between insolation forcing and monsoon response.

The East Asian monsoon is a subsystem of the Asian monsoon and forms one of the most extensive components of the Asia-Arabia-Australia monsoon, playing a key role in the transport of heat and moisture between latitudes (Ding 1994; Clift & Plumb 2008). Further, this redistribution of heat and moisture maintains the vast populations in China and SE Asia, and forms a fundamental component of global climate. A key challenge is to constrain the variation of the monsoon over millennial timescales, characterize any abrupt shifts in the system and consequently gain insight into the forcing mechanisms. Such information can be used to disentangle the tectonic and atmospheric/oceanic influences on the monsoon. One of the major obstacles to this is the relative paucity of available independent, radiometric dating methods that can be applied to monsoon archives at the temporal resolution required for the analysis of millennial scale events. In this study we present detailed summer monsoon reconstructions from wind-lain Chinese loess deposits, dated wholly by the optically stimulated luminescence (OSL) method (Aitken 1998) at high vertical sampling resolution (dating results presented in Stevens et al. 2006, 2008).

The results are from five sections across a transect of one of the major loess depocentres, the Chinese Loess Plateau, and complement recent reconstructions of the winter monsoon from the same sections, using the same independent OSL based age models (Stevens & Lu 2008). The results are used to assess the variation of the monsoon over the last c. 45 ka of the late Quaternary in the study region, and through comparisons to other independently dated records, to constrain forcing mechanisms and the relative timing of events. Only through independent dating can such comparisons be made. Prior to this a brief review of some recent attempts to date variations in the monsoon in China over the late Quaternary is presented, as well as an introduction to summer monsoon reconstruction from Chinese loess.

Chinese loess Wind-lain deposits of predominantly silt-sized Chinese loess and Red Clay provide a 22 Ma record of climate over part of the region affected by the monsoon (Liu & Ding 1998; Guo et al.

From: Clift, P. D., Tada, R. & Zheng, H. (eds) Monsoon Evolution and Tectonics –Climate Linkage in Asia. Geological Society, London, Special Publications, 342, 87– 108. DOI: 10.1144/SP342.8 0305-8719/10/$15.00 # The Geological Society of London 2010.

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2002). The thickest and most complete loess deposits occur in a region called the Chinese Loess Plateau, centred mainly on Shaanxi, Shanxi and Gansu Provinces (Fig. 1). Enhanced loess sedimentation occurs during periods of NW (winter) monsoon dominance with increased soil development during periods of strengthened SE (summer) monsoon. Most of the precipitation that falls on the Loess Plateau does so during the summer months as a result of penetration of the monsoon front into the region (Ding 1994). This monsoon precipitation causes pedogenic alteration of the loess parent material, including dissolution and reprecipitation of CaCO3, alteration of allogenic material to clay minerals, formation of superparamagnetic minerals, movement of mobile elements and associated colour changes (e.g. Liu & Ding 1998). As many of these processes are directly tied to the intensity of precipitation, an enhanced summer monsoon will ultimately result in greater pedogenesis. Where pedogenesis is particularly intense it is likely that bioturbation may result. Loess is widely regarded as the most detailed and complete terrestrial record of climate, comparable with marine sediments (Liu & Ding 1998; Porter 2001), and has been used extensively in the reconstruction of past monsoon variation (An et al. 2000; Ding et al. 2002; Guo et al. 2002). Summer monsoon reconstructions have been undertaken from a variety of proxies, including magnetic susceptibility (An et al. 1991; Maher 1998), reflectance/colour (Ji et al. 2001), stable isotopes (Rowe & Maher 2000), organic carbon and soil micromorphology (Kemp et al. 2001), Rb/Sr (Chen et al. 1999), CaCO3 (Shen et al. 1992), Chemical Weathering Index (Guo et al. 1996) and Fe-oxides (Sun & Ding 1998). By far the most widely used summer monsoon proxy has been magnetic susceptibility, despite debate over its ultimate origin (Zhou et al. 1990; Anderson & Hallet 1996; Porter et al. 2001; Maher et al. 2003; Balsam et al. 2004). In particular, the frequency dependent signal has been used widely as a proxy for summer monsoon induced pedogenesis, and is intended to estimate the relative contribution of fine, viscous grains at the border between superparamagnetic and single domain, to the total ferromagnetic assemblages (Chen et al. 1995; Dearing et al. 1996). Such grains are thought to be the products of repeated seasonal wetting and drying, as occurs in a monsoon climate (Maher & Thompson 1992). The errors on frequency dependence measurements are necessarily large and it has been suggested that high sensitivity, low frequency magnetic susceptibility (that measures the total ferromagnetic assemblage) may be a securer alternative (Stevens et al. 2007a, 2008). Despite these uncertainties, magnetic susceptibility measurements may provide one of the

best estimates of the relative intensity of summer monsoon induced pedogenesis. CaCO3 content has also been used as a proxy for past summer monsoon variation, as part of multi-proxy studies (Shen et al. 1992). It is generally assumed that allogenic primary CaCO3 is blown in from source regions and that during summer monsoon rainfall some of this is dissolved and removed by leaching processes. Below this active zone of dissolution the carbonate is then reprecipitated out in secondary form and, if this is particularly intense, may form a C horizon below the main soil. Much of the leached carbonate is likely to be precipitated out as nodules or micritic coatings (Kemp et al. 2001). Thus CaCO3 content is influenced both by the flux to the site, as well as summer monsoon intensity and is rather difficult to interpret in isolation. However, it can complement other summer monsoon proxy analyses and is useful indicator of pedogenetic processes. Of central importance in the reconstruction of monsoon changes over millennial and sub-orbital timescales is an accurate and precise age model. The majority of studies (e.g. Porter & An 1995; Ji et al. 2004) still employ age models derived through correlation to ‘known age’ sequences, such as the SPECMAP curve (Martinson et al. 1987), rather than through direct radiometric dating of the loess itself (e.g. Liu 1985; Buylaert et al. 2007). Largely, this has been due to the absence of a suitable technique, because radiocarbonbearing material is scarce and because U-series dates on carbonates are diagenetic rather than depositional ages (e.g. Rowe & Maher 2000). The widely applied correlation based age models are underpinned by a number of basic assumptions when used for constraining the timing of monsoon proxy variation: loess sedimentation proceeds essentially continuously; variations in the sedimentation rate are minor or can be accounted for through changes in grain-size; post depositional disturbance does not obscure syn-depositional climate records; and climate proxies or stratigraphic boundaries correlate with the SPECMAP curve. The most widely used models correlate visual or proxy stratigraphy in a type section with the marine record and use this association to mathematically constrain a grain-size sedimentation rate relationship that is then applied at other sections (Porter & An 1995; Nugteren et al. 2004). Many of the assumptions outlined above do not hold up for reconstructions on millennial timescales. Although reservations were expressed in the literature (e.g. Derbyshire et al. 1997), it has only recently been possible to test many of these assumptions using detailed OSL dating that directly constrains the age of the last exposure of loess sediment to light, and therefore provides a depositional

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Fig. 1. Topographic map of the study region indicating the sampled sites in this paper. Modern isohyets (Porter et al. 2001) and potential dust transport directions are shown as dashed lines and arrows respectively. The map of China (inset) shows the approximate location of the Chinese Loess Plateau (light shading), potential source regions (dark shading), major rivers and the study area (box).

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chronology (Wintle & Murray 2006; Stevens et al. 2007b). Many of these recent studies suggest that loess deposition cannot be considered as continuous (Singhvi et al. 2001; Stevens et al. 2006; Buylaert et al. 2007; Zhao et al. 2007; Jia et al. 2008), although there is some debate about this and not all detailed studies show up hiatuses (Roberts et al. 2001; Lai et al. 2007). There is an emerging consensus that loess deposition is more complex than previously thought, involving large changes in sedimentation rate that are not reflected in grainsize (Jia et al. 2007; Lai et al. 2007; Stevens et al. 2007a, 2008; Stevens & Lu 2008) and displaying evidence that proxies and stratigraphic boundaries in loess are not synchronous with marine oxygenisotope stage (MIS) records (Lai & Wintle 2006; Stevens et al. 2006). Many of these claims are reinforced by palaeomagnetic/pedogenetic evidence (Feng et al. 2004; Liu et al. 2004; Zhu et al. 2007). Stevens et al. (2006, 2007a) have also suggested that pedogenic disturbance at the base of the Holocene soil has, to varying degrees, removed any syn-depositional climate record at southern and central Loess Plateau sites for sediments at the base of the Holocene and upper Pleistocene. The combination of such effects will lead to errors in late Quaternary correlation based age models that may exceed 10 ka (Stevens et al. 2008). Clearly such errors render the investigation of millennial-scale climatic changes in loess impossible when using correlation based age models. As many prior findings from loess come from correlation based models there is now a pressing need to use age models derived from detailed radiometric dating for millennial scale climatic reconstructions from loess. Using the age models developed from over 100 OSL ages from five typical sections on the Loess Plateau (Fig. 1) presented in Stevens et al. (2006, 2008), this paper addresses this for summer monsoon proxies. Proxies for winter monsoon variation (grain-size and sedimentation rate) have also recently been examined using these age models (Stevens & Lu 2008).

Radiometric dating of the summer monsoon Beyond the results presented in Stevens et al. (2008), to our knowledge the only other radiometric dating based reconstruction of the summer monsoon from Chinese loess on a continuous time-axis has been that presented by Maher & Hu (2006). This study used the OSL dating results of Roberts et al. (2001) to develop a Holocene palaeo-precipitation reconstruction for East Asia based on magnetic susceptibility measurements. It was suggested that there were significant millennial-scale oscillations

in precipitation, some of which were in anti-phase with the SW Asian monsoon. Aside from loess, important reconstructions of the summer monsoon have come from speleothem records dated by 230Th (Wang et al. 2001, 2005, 2008; Yuan et al. 2004), 14 C dating of lake archives and ocean sediments (Dannenmann et al. 2003; Oppo & Sun 2005; Yancheva et al. 2007) and OSL dating of desert sediments (Lu et al. 2005). The reconstructions from speleothem records have been particularly significant, largely due to the extremely high precision and detailed dating potential offered by the technique. The findings have been used to suggest that the summer monsoon is affected by North Atlantic climatic variability and exhibits shifts on millennial timescales (Wang et al. 2001, 2005, 2008; Yuan et al. 2004). These studies also suggest that the monsoon is directly driven by summer insolation heating in the northern hemisphere over multi-millennial timescales and that Holocene changes in the monsoon are correlated to solar variation (Wang et al. 2005). Depending on conditions in the high-latitudes, this monsoon response to insolation may be both abrupt and delayed in comparison to deglacial melting (Cheng et al. 2006). However, just how much of a lag exists between insolation change and monsoon response may be dependent on contemporary boundary conditions and appears to be inconsistent between different points in time. Furthermore, there is still uncertainty as to how stable isotope records from speleothems can be translated into temperature and precipitation records (e.g. Kaufman & Dreybrodt 2004). Radiocarbon dating of oceanic sediments has been used to develop detailed late Quaternary monsoon reconstructions from the South China and Sulu Seas, as well as the broader Pacific Ocean (Stott et al. 2002; Dannenmann et al. 2003; Oppo & Sun 2005; Zhao et al. 2006). Changes in sea surface temperatures are suggested to be synchronous with events recorded in speleothem records (Wang et al. 2001), indicating a causal link, and one that extends on millennial timescales to the North Atlantic (Dannenmann et al. 2003). Reconstructions of monsoon change from radiocarbon dated lake sediments have been particularly useful in understanding the importance of rapid climatic changes during warm periods (Fang 1991; Bookhagen et al. 2005; Feng et al. 2005, 2007). The apparently anomalously wet MIS 3 [57–25 ka, as cited in Feng et al. (2007)] has been explained in terms of interplay between enhanced summer and winter insolation intensities and the effect of this on the summer monsoon (Feng et al. 2007), while in the Holocene, the concept of a widely applicable ‘Holocene optimum’ has been called into question. Further, Yancheva et al. (2007) have suggested that over the last 16 ka, the summer and winter

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monsoons are in direct antiphase due to both being forced via the position of the intertropical convergence zone. Recently, as described above, OSL dating has been applied to aeolian sediments in East Asia at unprecedented detail. The implications for understanding the monsoon have only just been investigated (Lu et al. 2005; Maher & Hu 2006; Stevens et al. 2008). Abrupt shifts in both the summer and winter monsoon have been postulated, although the precise relationship to the North Atlantic is unclear thus far. The summer monsoon may respond in a complex manner to a number of factors and appears to peak 2–4 thousand years after maximum insolation (Lu et al. 2005; Stevens et al. 2008). Stevens et al. (2008) suggested that this may be due to the distance of these records from the monsoon moisture source, and the consequent requirement of a sustained increase in forcing to enable the monsoon front to penetrate this region. Thus the possibility of temporal and geographical changes in the lag of the monsoon response to orbital forcing can now be investigated using diverse, independent, radiometrically dated records in a way not possible through correlation based approaches (which necessarily obscure information on leads and lags in the system).

Sites and methods The five sections at four sites investigated in this study, Shiguanzhai, Xunyi, Xifeng and two at Beiguoyuan, are located on a SE to NW transect across the Chinese Loess Plateau (site localities are shown in Fig. 1; basic unit stratigraphy is shown in Fig. 2) and are described in Stevens & Lu (2008). Modern annual precipitation varies across the sampled transect, causing significant pedogenic facies variation between the study sites. Beiguoyuan comprises two sections located near Huanxian, Gansu Province on the northern Loess Plateau (Fig. 1). The section with the most complete Holocene record (368370 21.300 N, 1078170 12.200 E, 1523 m a.s.l., hereafter termed Beiguoyuan Holocene section) lies several kilometres NW of the main section (368370 36.200 N, 1078160 57.400 E, 1545 m a.s.l., hereafter termed Beiguoyuan Main section) where a longer record of loess deposition is preserved. Sampling concentrated on the full Holocene section, entirely comprised of Black Loam (S0) and the upper 12.60 m of the main section, containing the Black Loam and upper Malan Loess (L1) Formations. The original section at Xifeng (Gansu Province), on the central Loess Plateau, is one of the most studied loess sections in China and has been used to develop age models, as well as records of long-term and rapid climate

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change (Kukla et al. 1988; Maher & Thompson 1991; Rousseau & Kukla 2000). A new Xifeng exposure (358320 09.400 N, 1078430 13.500 E, 1281 m a.s.l.) was sampled over the upper 4.30 m to include the entire Black Loam and upper Malan Loess. At Xunyi, Shaanxi Province (358010 25.100 N, 1088120 21.200 E, 1151 m a.s.l.) on the central Loess Plateau, strata were sampled over the upper 6.65 m to include the entire Black Loam and upper Malan Loess. At the Shiguanzhai site (348100 22.200 N, 1098110 45.500 E, 708 m a.s.l.), Shaanxi Province on the southeastern fringe of the plateau, the upper 1.75 m of the section was sampled, including the upper Malan Loess, Black Loam palaeosol and a distinct loess layer lying above, termed L0. Sampling for magnetic susceptibility analysis has been conducted twice at Beiguoyuan Main, Xifeng, Xunyi and Shiguanzhai sections, firstly at 5 cm intervals in 1995 and 1996, and secondly at 10 –20 cm intervals in 2003. Samples for OSL dating were taken at 10 –40 cm at all sites (Fig. 1). Full details of the OSL techniques tested and used in this study are presented in Stevens et al. (2007b) with age data presented in Stevens et al. (2006, 2008). A brief outline of the procedures used in age determination is presented here and OSL age versus depth data is plotted in Figure 3. Equivalent dose (De) values were measured using the SAR procedure (Murray & Wintle 2000) or using a standardized growth curve (Roberts & Duller 2004) performed on a Risø TL-DA-15 TL/ OSL reader. A blue LED (l ¼ 470 + 20 nm) stimulation source (60 s, c. 400 mJ cm21) was used on samples and the OSL signal was measured using a 9235QA photomultiplier tube filtered by 6 mm of Hoya U340 (Bøtter-Jensen et al. 2000). The signal was integrated from the first 0.6 s of stimulation minus a background estimated from the last 6 s of stimulation. All growth curves were fitted using a saturating exponential plus linear function (either sample-specific or a standardized growth curve). Aliquots yielding recycling ratios (Murray & Wintle 2000) or IR ratios (Duller 2003) differing from unity by greater than 10% were rejected. The uncertainty on individual De values was estimated using Monte Carlo simulation and a weighted mean De (with one standard error uncertainty) was calculated for each sample (typically .12 aliquots). Dose rates were calculated using 238U, 232Th and 40K contents measured using Inductively Coupled Plasma Mass Spectrometry (ICP-MS) at Kingston University School of Earth Sciences and Geography (Agilent 7500) and Atomic Emission Spectrometry (ICP-AES) at Royal Holloway University of London, Department of Earth Sciences (Perkin Elmer Optima). Age uncertainties are based on the propagation, in quadrature, of individual errors for all measured quantities, which if

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Fig. 2. Visual stratigraphic logs of the studied sections, including the nomenclature for stratigraphic units. Shading and bracketed Munsell values are for general guidance only. Visual stratigraphy varies more than is represented by the Munsell values. There is a general darkening in colour from north to south in studied sections (Fig. 1) and, although overall Munsell values remain similar for corresponding units between sections, there are darker excursions within strata in the south and lighter excursions within strata to the north.

RADIOMETRIC DATING OF THE MONSOON Fig. 3. OSL age (+1 standard error) versus depth for all the study sites. Full age data are presented in Stevens et al. (2006, 2008) and rationale for age exclusion/inclusion in the age models presented therein.

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unknown are taken as 10%. The cosmic dose was calculated using present day burial depth (Prescott & Hutton 1994). Full details of the derivation of depth-continuous age models from individual OSL ages are presented in Stevens et al. (2008). The age models extend to c. 15, 40, 29, 46 and 30 ka for Beiguoyuan Holocene, Beiguoyuan Main, Xifeng, Xunyi and Shiguanzhai sections respectively. Such age models allow analysis of magnetic susceptibility, sedimentation rate and CaCO3 content on a timecontinuous age axis that takes account of hiatuses and disturbance in the record, as well as changes in sedimentation rate, whilst being completely independent of tuning to orbital time series. Calculation of sedimentation rates is presented in Stevens & Lu (2008). Samples for magnetic susceptibility analysis were pre-treated by air drying, disaggregation and sieving at 2 mm prior to weighing and measurement in 25  25 mm diamagnetic pots. Measurements were taken using high (4.7 kHz) and low (0.47 kHz) frequency fields at high sensitivity (0.1) using a Bartington MS2 magnetic susceptibility meter at Oxford University Centre for the Environment and the Chinese Academy of Sciences, Xi’an. Bulk low frequency susceptibility (xlf) measurements are given in units of m3 kg21 and % frequency dependence (xfd) is defined as

Beiguoyuan Holocene section The Holocene section at Beiguoyuan shows a general trend of decreasing magnetic susceptibility values and increasing CaCO3 content with depth (Fig. 4). xfd and xlf show similar trends with

xlf  xhf xlf for 2003 data only (where xhf is bulk high frequency susceptibility). Frequency dependence is intended to estimate the relative contribution of fine viscous grains at the border between superparamagnetic and single domain to the total ferromagnetic assemblage (Dearing et al. 1996). Magnetic susceptibility changes may correspond to changes in summer monsoon induced weathering with the single domain and total ferromagnetic assemblage (as represented by xfd) being the most sensitive indicator of the East Asian SE monsoon (Maher & Thompson 1991). However, due to the nature of the calculation involved, propagated errors are large for xfd and on average lie at about 20%, increasing in palaeosol units (Stevens et al. 2007a). xlf is used in this study to represent the total ferromagnetic assemblage and errors are generally of the order of ,1%. CaCO3 content was estimated through weighing samples before and after treatment with concentrated HCl.

Results The results are presented first by depth, and then by OSL age for individual sites.

Fig. 4. Proxy variations with depth at Beiguoyuan Holocene section. (a) Variations of xfd and CaCO3 (%) with depth; (b) xlf (m3kg21 – 2003 data) with depth. The whole section is the Black Loam Formation (S0).

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depth. Maximum soil development therefore appears to be in the upper part of the profile, with secondary enrichment of diagenetic CaCO3 occurring lower in the sequence. Figure 5 shows the same data plotted by OSL age, alongside calculated sedimentation rate. Peak magnetic susceptibility values occur during the middle part of the Holocene (i.e. ,5 ka), although the later Holocene is missing and an unconformity occurs in latest glacial times. The secondary reprecipitation of CaCO3 at the base of the Holocene supports the contention that peak soil development occurred in the middle part of the Holocene. The sedimentation rate also increases during the middle Holocene, to similar levels to that during the late glacial.

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Beiguoyuan main section The main section at Beiguoyuan shows similar patterns in both magnetic susceptibility and CaCO3 content with depth (Fig. 6). Furthermore, xlf collected in both 1996 and 2003 shows similar behaviour, except at the 5 cm scale and in the upper c. 1.20 m. Values of xlf decrease above 12.00 m depth, increase again between 10.00 and 8.00 m, and remain steady until the Black Loam, where values climb significantly. Figure 7 shows the same data plotted by OSL age, alongside calculated sedimentation rate. Values of xlf drop down from a peak at about 37 ka, and subsequently increase to similar levels at 29 ka. At 25 ka xlf drops to lower levels

Fig. 5. Proxy variations with age at Beiguoyuan Holocene section. (a) Variations of xfd and CaCO3 (%) with age; (b) xlf (m3kg21 – 2003 data) and sedimentation rate (cm ka21) with age.

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the lower part of the section (above 2.00 m), and includes some high frequency variations (Fig. 8a). The xlf data collected in 1996 broadly match, excepting the uppermost 1.50 m (Fig. 8b). Once again, the peaks in magnetic susceptibility are best expressed in the xlf data, especially when the errors are considered. Values are also higher towards the base of the sequence. CaCO3 values show high magnitude, abrupt shifts and generally decreases up-section. The upper 1.00 m shows low CaCO3 values that can be attributed to leaching during pedogenesis. Figure 9 shows the same data plotted by OSL age, alongside calculated sedimentation rate. The clear peak in magnetic susceptibility data is shown in the mid/late Holocene (6–2 ka), although whether the early Holocene has similar values is uncertain because of bioturbation. Between 30 and 20 ka, magnetic susceptibility values decreased slowly, although there is a sharp drop in xlf values at 27 ka. CaCO3 values show similar trends between 30 and 20 ka, but no slow change to lower values. In the Holocene, CaCO3 shifts are in direct antiphase with xfd. Sedimentation rates show broadly similar patterns to those of xfd, although with highest values around 29 ka.

Xunyi

Fig. 6. Proxy variations with depth at Beiguoyuan Main section. (a) Variations of xfd and CaCO3 (%) with depth; (b) xlf (m3kg21 – 2003 and 1996 data) with depth.

again, where it fluctuates until rising again to peak Holocene values just after 10 ka and an unconformity in the record (Stevens et al. 2006). Sedimentation rates show three pronounced peaks centred on 29, 24 and 20 ka. The less extensive dataset for CaCO3 and xfd shows more variation than xlf (Fig. 7a). The two measurements show similar trends and appear to show earlier, but less rapid increases associated with the transition into the Holocene.

Xifeng The section at Xifeng shows higher magnetic susceptibility values in the upper part compared to

xfd values show minimal percentage variations. When the c. 20% errors are also considered this limits description of any significant shifts (Fig. 10a). Nevertheless, it seems to broadly parallel changes in xlf, which generally matches between 2003 and 1996 data (Fig. 10b), despite some of the finer scale changes being slightly offset. CaCO3 shows very significant, high frequency shifts with depth (Fig. 10a). Values are slightly reduced in the uppermost part of the sequence. Figure 11 shows the same data plotted by OSL age, alongside calculated sedimentation rate. Values of magnetic susceptibility are highest in the Holocene, although xfd values show less pronounced low values in older sediments. From 45 to 20 ka there is an intermittent lowering of xlf values at the site. Changes in xfd over this period are to some extent repeated by changes in CaCO3 values, although they are in antiphase during the Holocene, indicating leaching in this part of the sequence. Sedimentation rates show three clear peaks, at 45, 38, 23 ka, and are reduced in the Holocene.

Shiguanzhai As at Xunyi, xfd values show very low percentage variations (Fig. 12a). In combination with the errors this limits the use of this proxy at the site. CaCO3 shows far larger percentage variations, with a general drop in values up-sequence,

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Fig. 7. Proxy variations with age at Beiguoyuan Main section. (a) Variations of xfd and CaCO3 (%) with age; (b) xlf (m3kg21 – 1996 data) and sedimentation rate (cm ka21) with age.

associated with leaching in the Black Loam Formation. Changes in xlf values are similar between the two datasets, although slightly offset between 0.50 and 0.80 m. Figure 13 shows the same data plotted by OSL age, alongside calculated sedimentation rates. The large gap in the record is likely due to bioturbation (Stevens et al. 2006). Magnetic susceptibility values show similar trends with age, although xlf shows a later Holocene peak. Both show lower values from 30 to 23 ka. CaCO3 shows shifts of the opposite sign to xfd, with highest values between 30 and 23 ka. The sedimentation rate also exhibits this pattern, with lowest values during the Holocene.

Discussion From the data presented above some basic conclusions can be drawn. The two measures of magnetic susceptibility (xfd and xlf) show similar broad trends. At finer scales (sub-millennial and ,20 cm) the measures exhibit differences. Whether this is a consequence of different controls on individual proxy development or the large error term on xfd is unclear. Interpretation of xfd data at more southern and central sites (where xfd values are higher) is limited by the very large errors (values of xfd are higher, leading to larger absolute errors). For this reason the xlf variations may be a

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Fig. 8. Proxy variations with depth at Xifeng. (a) Variations of xfd and CaCO3 (%) with depth; (b) xlf (m3 kg21 – 2003 and 1996 data) with depth.

safer measure of pedogenesis and consequently, summer monsoon development. Sedimentation rate may have had an influence on the development of pedogenesis proxies through dilution of pedogenetic signatures and changing of

the length of exposure of sediments to the weathering zone (Anderson & Hallet 1996). The findings presented here suggest that there is very little similarity in the relationship between sedimentation rate and magnetic susceptibility across the sites. For example, at Xifeng the two proxies change in broad phase and in the same sign (Fig. 9b), indicating that when pedogenesis increases, sedimentation rates also increased. Roberts et al. (2001) reported that magnetic susceptibility variations appeared independent of OSL derived sedimentation rate changes, implying that precipitation was the dominant control on magnetic susceptibility and that dilution of the signal by changing accumulation rates is negligible (Maher et al. 2003). This further implies that summer and winter monsoons do not change in direct antiphase over this timescale [contrasting with the lake record of Yancheva et al. (2007)] and that the magnetic susceptibility signal develops rapidly. In contrast to Xifeng, at Shiguanzhai the two measures are in antiphase (Fig. 13b). This contrast between sites reinforces the suggestion that sedimentation rate may be a complex response to a wide range of factors, limiting its use as a winter monsoon proxy (Stevens & Lu 2008). It also demonstrates that it is possible to have increased summer monsoon precipitation whilst conditions for the formation, transport and deposition of dust are conducive to enhanced loess accumulation. Such conditions are normally associated with at least moderate winter monsoon activity. Thus, on the Loess Plateau at least, the two monsoon systems appear not to vary in complete antiphase and that episodes of enhanced winter and summer monsoon occurring at the same time are possible. The contrasting relationship between sites also supports the assertion that the dominant control on magnetic susceptibility is precipitation in the region (Roberts et al. 2001; Maher et al. 2003). Another consideration is that while the two xlf datasets are in general agreement, on the finer ,20 cm scale, there is some disparity. This appears to be most significant in the Black Loam Formation. The cause of these discrepancies may be sampling error, instrumental variation, machine error, lateral changes in pedogenesis at sites, or a combination. In any case this suggests that when sampling is conducted at ,20 cm intervals caution should be exercised when ascribing such fine-scale variation to broad scale monsoon changes. That these discrepancies seem to be especially manifest in the Holocene Black Loam Formation (with the most intense pedogenesis) reinforces the possibility that lateral changes in bioturbation may be an important factor in the fine scale differences, and that this will affect the preserved record in loess soils in particular.

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Fig. 9. Proxy variations with age at Xifeng. (a) Variations of xfd and CaCO3 (%) with age; (b) xlf (m3 kg21 – 1996 data) and sedimentation rate (cm ka21) with age.

CaCO3 changes show a very complex signal with many high frequency variations. This might suggest that it is particularly useful for analysis of sub-millennial scale variation in the monsoon. Carbonate content will be influenced by in-situ carbonate leaching and precipitation as well as (presumably) allochthonous primary carbonate sedimentation. This dual influence is evident in the fact that changes do not consistently correspond with magnetic susceptibility variation and ensures that use of CaCO3 variations for climatic interpretation on its own is problematic. In order to analyse variations in the broad-scale summer monsoon using the proxies it is necessary to examine data from all the sites plotted by OSL

age. Without recourse to correlating sequences based on stratigraphy this is only possible using detailed independent dating. Thus, real differences in the timing of monsoon changes between sites can be examined; the first time such a comparison has been made for summer monsoon data for loess. Figure 14 shows variations in xlf and CaCO3 by OSL age for all the study sites. xlf variations are broadly similar between sites (Fig. 14a), reinforcing the signal’s use as a monsoon proxy. There are some discrepancies in the sub-millennial scale variation. In particular, in the sub-millennial patterns during MIS 3 and the onset of the Holocene. In general, values are highest in the Holocene (12 –0 ka), followed by MIS 3 (59 –24 ka), and

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Fig. 10. Proxy variations with depth at Xunyi. (a) Variations of xfd and CaCO3 (%) with depth; (b) xlf (m3 kg21 – 2003 and 1996 data) with depth.

lowest in MIS 2 [24 –12 ka; all ages as in Martinson et al. (1987)]. In contrast to other sites, Beiguoyuan Main shows some variation in MIS 3 (Stevens et al. 2008). It is uncertain what the cause of this variation is and it may be related to the fact that Beiguoyuan Main lies at the northernmost limits of current monsoon influence and will therefore be

very sensitive to small changes in monsoon intensity and front penetration. The transitions between these peaks and troughs of monsoon intensity are relatively rapid, again potentially because of the location of Beiguoyuan Main. Further to the south these rapid transitions are not evident. During the last glacial there appear to be few rapid changes at any site, although the record is interrupted by hiatuses. The transition to the Holocene is rapid in all sites, especially at Beiguoyuan Main, indicating a sudden increase in weathering, caused by significant intensification of the monsoon. Thus, precipitation on the Loess Plateau appears to have rapidly increased just after Holocene onset (9–8 ka). Identifying a peak in weathering is difficult because of the fragmentary record, but does appear to occur in the mid Holocene (between 4 and 9 ka). However, some records are missing peak weathering values. A midHolocene timing for the ‘Holocene optimum’ is slightly later (up to 5 ka) than has previously been recognized for the region (An et al. 2000; He et al. 2004). The records of CaCO3 are far more difficult to interpret because the variations are less consistent between sites (Fig. 14b). Nonetheless, some trends can be identified for all sites. Most obviously all sites show significant decalcification during the Holocene, which we interpret to be caused by leaching. As would be expected, this is most severe for the southern and central sites. During the last glacial period average values remain fairly constant for the sites, although there is lots of variation about the mean and very little of this is coincident between the sites. Beyond the latest episode of carbonate leaching due to enhanced summer monsoon conditions there is little environmental interpretation possible. The high frequency variation is likely a complex response to changes in leaching, re-precipitation and influx of allochthonous carbonate. When the CaCO3 and xlf values are normalized by the average value (Fig. 15) the trends between sites can be more readily identified. It also enables comparison with other independently dated climate reconstructions and timescales. Changes in speleothem d18O from East/Central China (Wang et al. 2001, 2008) show broadly similar variations to xlf but much of the millennial scale variation in d18O is not replicated in the magnetic susceptibility records (Fig. 15a). The rise towards peak Holocene values in d18O appears to precede the rises in xlf by a few thousand years. Stevens et al. (2008) explained this for Beiguoyuan Main in terms of the geographic distance between the records and the distance of the Loess Plateau from the monsoon moisture source. It could be expected that only at maximum insolation forcing would a measurable

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Fig. 11. Proxy variations with age at Xunyi. (a) Variations of xfd and CaCO3 (%) with age; (b) xlf (m3kg21 – 1996 data) and sedimentation rate (cm ka21) with age.

response occur in terms of rainfall on the Loess Plateau because the region lies at the limit of recent monsoon front penetration. Another potential reason for the difference lies in the incompleteness of the loess records (Fig. 15). At most sites the record of the late glacial –early Holocene is missing, possibly due to either erosion or enhanced pedogenesis. It is feasible that rainfall increased in the region during that time but the record is not present today. Another difference between the loess and speleothem records is the lack of an anomalous reduction in summer monsoon strength in the

speleothem record during MIS 3 (59 –24 ka), in contrast to that suggested at Beiguoyuan Main. A potential explanation lies in the winter monsoon’s strength. At this point winter monsoon strength was relatively high (Stevens & Lu 2008) and this may have affected Beiguoyuan Main in particular, due to its proximity to the arid continental interior and the Siberian High pressure system. This high pressure zone may have reduced the penetration or influence of the summer monsoon in this region, or delayed its onset. Despite these differences, both records do record a significant drop in summer monsoon intensity at around 25 ka. When the

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Fig. 12. Proxy variations with depth at Shiguanzhai. (a) Variations of xfd and CaCO3 (%) with depth; (b) xlf (m3kg21 – 2003 and 1996 data) with depth.

speleothem record is compared to changes in CaCO3 (Fig. 15d) some broadly similar trends are also apparent. However, the millennial scale variation between the records bears little correspondence.

When compared to changes in June insolation at 308N (Berger & Loutre 1991), xlf changes at the study sites show a lack of correspondence (Fig. 15b). This may also be interpreted as a lag of monsoon response to insolation forcing, and this is consistent with the Loess Plateau’s location far from the monsoon moisture sources (Stevens et al. 2008). A significant increase in forcing is required to drive precipitation increases in remote locations, like the Loess Plateau, and only after prolonged intensification of forcing would the monsoon front penetrate these regions and force a consummate change in summer monsoon proxies. Even if this hypothesis is accepted there are still significantly different patterns in insolation forcing and response. For example, the anomalous drop in magnetic susceptibility values between 38 and 30 ka at Beiguoyuan Main. At the other sites there is little variation in the monsoon proxies between 45 and 25 ka, beyond a gradual decrease in values, despite significant insolation changes. To some extent, this is also a characteristic of the speleothem record and this variation was still interpreted as insolation driven. Changes in CaCO3 (Fig. 15e) show even less correspondence with insolation. Finally, if variations in xlf are compared to the global record of ice volume changes from composites of stacked d18O records (Martinson et al. 1987) then arguably the best correspondence between the data is achieved (Fig. 15c). The shift in xlf between the Pleistocene and the Holocene is similar in rate to the SPECMAP record and the slow decline in normalized d18O from 50 –18 ka also matches the xlf record at the study sites. After 18 ka, the SPECMAP curve normalized d18O values increase progressively until maximum Holocene values, reached at about 9 ka. This contrasts with the xlf record that shows more abrupt and delayed increases, taking place from 9 ka (although this interpretation is limited by the incomplete record preserved). Variations in CaCO3 are also most closely approximated by the oceanic d18O record of ice volume, although the records are still quite different (Fig. 15f). The key uncertainty is whether this association between ice volume and summer monsoon proxies implies causality, or whether both changes are ultimately modulated by changes in other forcing mechanisms with similar characteristics. The summer monsoon is believed to be forced by changes in insolation, via direct heating of the western Pacific (e.g. Wang et al. 2008). This would imply that any correspondence between ice volume changes and summer monsoon intensity from loess proxies would be due to them having the same ultimate forcing mechanism; insolation. It seems feasible that changes in ice volume will have an influence on monsoon source region sea surface temperature through changes in sea level and ocean circulation. Further, changes in ice

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Fig. 13. Proxy variations with age at Shiguanzhai. (a) Variations of xfd and CaCO3 (%) with age; (b) xlf (m3kg21 – 1996 data) and sedimentation rate (cm ka21) with age.

volume will alter wind fields and circulation systems (such as the Siberian High pressure system and westerly flow) that would likely affect the penetration of the summer monsoon system into the arid continental interior of Asia. Changes in winter monsoon proxies from the Loess Plateau also appear to show similarities with global ice volume changes (Stevens & Lu 2008) and whereas summer and winter monsoon records are not directly correlated at millennial timescales, they do show broad multi-millennial timescale antiphase behaviour (Stevens et al. 2008). Another complicating factor is the lag of summer monsoon increases behind both insolation, and to a lesser extent, ice volume. This suggests that further factors are involved, such as the geographic locality (see

above) and that it is the combined influence of this and a strong winter monsoon that delays the onset of an enhanced summer monsoon in the Holocene. This possibility is outlined below. Potentially, increases in the strength of the Siberian High pressure system [which would likely be influenced by ice volume, as it is influenced by snow depth (Jhun & Lee 2004)] would limit the extent to which the summer monsoon can penetrate the continental interior by reducing the pressure/temperature/humidity contrast between continent and ocean year round [this contrast is ultimately what drives summer monsoon circulation (Clift & Plumb 2008)]. At the remote ends of the penetration of the summer monsoon (i.e. the Loess Plateau) this would have been felt through

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Fig. 14. Variations in (a) xlf (m3 kg21 – 1996 data, except Beiguoyuan Holocene: 2003 data); and (b) CaCO3 (%) with age for all 5 sample sites. Key: BH, Beiguoyuan Holocene; BM, Beiguoyuan Main; XF, Xifeng; XY, Xunyi; SG, Shiguanzhai. Vertical black lines delineate approximate MIS boundaries (Martinson et al. 1987).

significantly reduced rainfall as the monsoon front fails to penetrate the region. Thus, the ice volume might affect the strength of the summer monsoon, and at extreme ends of the monsoon front’s penetration, this may be the dominant factor affecting rainfall. This may not be the case at more moisture source proximal locations where conditions in the deep continental interior will have less of an effect. Thus during the termination of the last glacial, while insolation was still increasing and the ice sheets were still melting, enhanced summer monsoonal precipitation may not have been experienced over the Loess Plateau, despite increases in forcing. Enhanced precipitation may have taken place closer to the source region but a combination of the prevailing influence of the Siberian High in the continental interior and the slow progression

of the summer monsoon front into the region, penetrating only at peak insolation forcing (9– 10 ka; Stevens et al. 2008), will have limited monsoonal precipitation on the Loess Plateau. Several hundred kilometres to the south at Sanbao cave in Central China (Wang et al. 2008), precipitation records also appear to lag insolation changes, although by less than for the loess records presented here. This hypothesis of the influence of ice volume and extent of monsoon front penetration affecting regional manifestations of monsoon driven precipitation can be directly tested by using independent dating to obtain summer monsoon records from locations throughout the range of the summer monsoon in Asia. A further complicating factor may be sedimentation rate. Sedimentation rate varies considerably

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Fig. 15. Normalized (to average value) variations in (a), (b), (c) xlf (m3 kg21; 1996 data, except Beiguoyuan Holocene: 2003 data) and (d), (e), (f) CaCO3 (%) with OSL age for all five sample sites. Data are plotted alongside (black lines): (a), (d) d18O (VPDB) from stalagmites in Sanbao and Hulu caves, China (Sanbao cave values plotted 1.6 per mil more positive to account for lower Sanbao cave values; Wang et al. 2001, 2008); (b), (e) June insolation at 308N (Wm2) (Berger & Loutre 1991) and; (c), (f) SPECMAP d18O (normalized) (Martinson et al. 1987). Line colours for individual sites are the same as for Figure 14. Vertical black lines delineate approximate MIS boundaries (Martinson et al. 1987).

between sites and over time on the Loess Plateau (Stevens & Lu 2008). These changes still have the potential to modulate the way that summer monsoon proxies are recorded in loess, despite the lack of relationship between sedimentation rate and magnetic susceptibility. It is possible that any direct link between insolation, moisture source temperature and summer monsoon proxies is obscured by this factor.

Conclusions Changes in summer monsoon proxies (CaCO3 and magnetic susceptibility) have been analysed both by depth and by OSL age for five sites that form a

transect across the Chinese Loess Plateau. The use of wholly independent, radiometric age models allows new insight into the forcing mechanisms behind monsoon change in this region, close to the limit of the summer monsoons influence. Changes in CaCO3 are complicated by the multiple influences on its variation. However, changes in magnetic susceptibility proxies (xlf and xfd) are broadly similar and although xfd has been used in the past as a proxy for soil development, the large errors are prohibitive for more southern sites. xlf shows significant variation and can be used to show changes in summer monsoon intensity, despite the uncertain influence of sedimentation rate on the proxy.

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Changes in xlf show that the summer monsoon in this part of China does not exhibit many high frequency shifts (although this is possibly due to the nature of signal acquisition) and show general decreasing trends from c. 50–18 ka. Early in the Holocene xlf values increase dramatically at the study sites and show generally good correspondence. The Holocene ‘optimum’ in the region may be younger than previously suggested, with peak pedogenesis occurring at 9–4 ka, rather than before 9 ka (e.g. An et al. 2000). Indeed, the hypothesis that the rainfall peak in China during the Holocene gets younger progressively to the SE (An et al. 2000) contrasts with the independently dated records from speleothems (Wang et al. 2001) and the loess records presented here. Changes during MIS 3 (59–24 ka) at the northernmost site may be forced by small changes in summer monsoon strength that force a non-linear response in precipitation strength in the region, due to the site’s position at the furthest extent of the monsoons influence. When compared to independently dated records of summer monsoon variation closer to the moisture source, insolation and global ice volume, the closest correspondence is to ice volume. This is hypothesized to be due either to forcing via ice volume controlled atmospheric and oceanic circulation (strength of the Siberian High), or due to similar responses to insolation forcing. Despite the latter being more widely suggested previously, the former hypothesis is suggested to be viable and testable. Key is the use of independently dated records using radiometric techniques at very high sampling resolution. Only with such data can enquiries into monsoon forcing be made without recourse to correlation techniques and consequent circular logic. Support from the Nordic Centre of Excellence programme of the Joint Committee of the Nordic Natural Science Research Councils is gratefully acknowledged. The authors thank Yi Shuangwen and Sun Xuefeng for help in the field, NERC for access to their ICP-MS and -AES facilities (Awards OSS/279/0205 and OSS/302/1105) and the 211, 985 and 111 program of Nanjing University for fieldwork support. T.S. thanks Jesus College, Oxford University Centre for the Environment, the Royal Society and the Geological Society of London for financial support. Thanks to H. Lunn for CaCO3 and magnetic susceptibility work and M. Smith and C. Ivison for help with Figure 1.

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Climate shift recorded at around 10 Ma in Miocene succession of Samburu Hills, northern Kenya Rift, and its significance TETSUYA SAKAI1*, MOTOTAKA SANEYOSHI2, SATOSHI TANAKA3, YOSHIHIRO SAWADA1, MASATO NAKATSUKASA4, EMMA MBUA5 & HIDEMI ISHIDA6 1

Department of Geoscience, Shimane University, Matsue 690-8504, Japan 2

Centre for Paleobiological Research, Biochemical Laboratories Inc., Okayama 700-0907, Japan

3

Faculty of Education, Kyoto University of Education, Kyoto 612-8522, Japan 4

Department of Zoology, Graduate School of Science, Kyoto University, Kyoto 606-8502, Japan 5

Department of Earth Sciences, National Museums of Kenya, P.O. Box 40658-00100, Nairobi, Kenya

6

School of Human Nursing, University of Shiga Prefecture, Hikone, Shiga 522-8533, Japan *Corresponding author (e-mail: [email protected]) Abstract: A significant climate shift around 9.6 Ma has been detected from the Middle to Upper Miocene Aka Aiteputh and Namurungule Formations exposed in the Samburu Hills, northern Kenya. Around 9.6 Ma, changes in sediments are recorded from the red soil-dominated interval of the upper Aka Aiteputh Formation to the lacustrine and deltaic facies of the lower Namurungule Formation, containing open woodland/savanna mammalian fauna. These reveal a shift from a dry climate with seasonal precipitation to a climate with strong seasonality. In particular, an increase in precipitation was recorded by the predominance of lacustrine facies. This shift happened at around the same time as the intensification of the Indian summer monsoon that has been detected in the Himalayas and some of surrounding regions. There are two scenarios that could explain the increased precipitation at the beginning of the deposition of the Namurungule Formation: (1) enhanced moisture transport by the Intertropical Convergence Zone (ITCZ), synchronized with Indian summer monsoon intensification, or (2) intensification of the Indian summer monsoon itself, permitting moisture to penetrate deep into East Africa if the altitude of the rifted area was lower than it is now. Presently, the former is considered to be the more plausible explanation for the climate shift detected in the Samburu Hills.

Although reconstructing East Africa’s environment is crucial for understanding the evolution of hominoids and hominids (Brown et al. 1985; White et al. 1994; Ishida & Pickford 1997; Senut et al. 2001), an understanding of the period just before the first hominids appeared, between 10 –7 Ma, is lacking because of the scarcity of exposed sedimentary successions in the region. The time period in which hominoids and hominids evolved corresponds to a time of major global climate shift (e.g. Trauth et al. 2005; Se´galen et al. 2007). The period from 12 –8 Ma has been proposed as a time when the Indian summer monsoon started and intensified (e.g. Clemens 2006). Climate changes have also been documented from terrestrial sediment successions in the Siwalik Hills, the frontal region of the Himalayas. Dettman

et al. (2001) analysed O and C isotopes of fossil molluscan shells and revealed larger variations of the isotope ratios within a single molluscan shell body from several horizons after 10.7 Ma, suggestive of enhanced seasonality (with a distinct wet and dry season), and increasing aridity at 7.5 Ma. The oldest visible sediment succession recorded is from 10.5 Ma. Nakayama & Ulak (1999) and Ulak & Nakayama (2001) detected changes in fluvial facies between 10.5 to 9.5 Ma from the Bakiya Khola, Surai Khola and Tinau Khola sections in Nepal, resulting from an increase in precipitation. Other indicators such as plant and mammalian fossils also show evidence of climate changes around 12–10 Ma in the Himalayan frontal region as well as in the surrounding areas, such as the Tibetan Plateau (Barry et al. 1995;

From: Clift, P. D., Tada, R. & Zheng, H. (eds) Monsoon Evolution and Tectonics –Climate Linkage in Asia. Geological Society, London, Special Publications, 342, 109–127. DOI: 10.1144/SP342.9 0305-8719/10/$15.00 # The Geological Society of London 2010.

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Konomatsu 1997; Dettman et al. 2003) before the major climate shift around 9– 7 Ma (Quade et al. 1989). Although there have been many studies regarding the climate in the Himalayas, the Tibetan Plateau, and the surrounding Asian regions, there have been few climatological studies in East Africa covering the Miocene period, although the area is located along the summer wind pathway to the Himalayas (e.g. Camberlin 1996; Gatebe et al. 1999). The purpose of this study is to characterize the sediments from around 10 Ma that are exposed in the Samburu Hills, in the northern Kenya Rift, in order to decipher the climate shift from the succession and to discuss the causes of the shift.

Geological setting The Samburu Hills are located in the northern part of Kenya, c. 50 km south of Lake Turkana (Fig. 1). The Samburu Hills [altitude ranging from 1400–400 m above sea-level (a.s.l.)] form the shoulder of the eastern flank of the Kenya Rift. The rift centre, the Suguta Valley (c. 400 m a.s.l.), extends north– south and runs west of the hills (Fig. 1). The onset of volcanism in the northern part of Kenya occurred at around 30 Ma (e.g. Baker 1986) and its activity propagated towards the south. Basin individualization and uplift of the rift basin shoulders occurred between 11–5.3 Ma (Chorowicz 2005). Volcanism and sediment accumulation started around 20 Ma in the Samburu Hills (Sawada et al. 2006). From bottom to top, the Miocene succession in the Samburu Hills (Figs 2 & 3) consists of the Nachola Formation, which covers the basement rock (mainly gneiss and granitic rocks of the Mozambique Belt), the Aka Aiteputh, the Namurungule, and the Kongia Formations (Figs 2 & 3) (Sawada et al. 1987, 2006). At the top of the Samburu Hills succession is the Pleistocene Tirr Tirr Formation (Fig. 3). The exposed succession tends to young towards the present rift centre (Suguta Valley) due to the successive western shift of the flexure zone (Willamas & Chapman 1986). The Nachola Formation comprises mainly layers of basalt and trachytic lava, pyroclastic rocks, and local siliciclastic sedimentary rocks. The lava layers within the formation have been dated as between 19.2 and 15.0 Ma using K –Ar dating [see Sawada et al. (1998) for details]. The Aka Aiteputh Formation (15.0 –9.6 Ma) conformably covers the Nachola Formation and the lower and middle parts of the Aka Aiteputh Formation consist mainly of basaltic lava and red soil beds with calcrete layers. The base of the upper Aka Aiteputh Formation is marked by a minor unconformity (Figs 3 & 4). The minor unconformity surface is mostly planar and concave-up near the north–south trending fault in

the western Samburu Hills. It is associated with basin fragmentation and the formation of a series of half grabens. The fragmented basins are visually represented as the scattered Namurungule Formation on the geological map (Fig. 2). The upper part of the Aka Aiteputh Formation comprises predominantly red soil beds (with common calcrete layers) and basalt lava layers. Some gravelly sandstone beds are intercalated within the red soil beds. The basal part of the Namurungule Formation comprises conglomerate and sandstone beds forming a small alluvial fan below alternating sandstones and mudstones originating from a lacustrine delta (Figs 4 & 5) (Saneyoshi et al. 2006). Up to 20 m of poorly sorted tuffaceous mudstone beds with abundant gravel (from granule to boulder size) of lahar origin is interbedded in the middle part of the formation. The base of the formation has been dated at 9.6 Ma (Sawada et al. 1998) and the horizon of the shift from reverse to normal palaeomagnetic polarity in the upper part has been dated to 9.3 Ma (Sawada et al. 2006). The Namurungule Formation is unconformably overlain by the Kongia Formation, which is represented by basalt lavas and red soil beds. The Kongia Formation deeply truncates the more inclined Aka Aiteputh and Namurungule Formations owing to a tectonic event prior to deposition of the Kongia Formation (see Sawada et al. 2006). This formation was dated to 7.3–5.3 Ma by K –Ar methods (Itaya & Sawada 1987; Tatsumi & Kimura 1991; Sawada et al. 1998). In the northwestern part of the Samburu Hills, the Tirr Tirr Formation (4.1– 3.6 Ma) comprises basalt and trachyte lavas (Kabeto et al. 2001) and unconformably covers the Kongia Formation (Fig. 2). This sequence has a maximum thickness of about 180 m.

Depositional facies of the upper Aka Aiteputh and Namurungule formations Upper Aka Aiteputh Formation The base of the upper Aka Aiteputh Formation (c. 10.1 Ma; Fig. 4) is represented by conglomerate beds (up to 10 m thick). The conglomerate beds are generally lenticular in shape, filling the base of small half grabens. The upper part of the formation above the conglomerate beds contains five distinct basalt lava layers (Lava 1– 5: each 3 –10 m thick) and intervening red soil beds. There are also minor lava layers. The red soil beds, about 5 m in thickness, contain abundant calcrete layers (Fig. 6a), caliches and remains of rootlets. Thin calcrete beds (0.01– 1 m thick) are commonly interbedded within the red soil beds, and they contain moulds of roots (up to 5 mm in diameter) (Fig. 6b) and laminations. Up to 1 m of white chert beds,

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Fig. 1. Location and geological map of the Samburu Hills. (a) Location of study area; (b) topographic and simplified geological structures of northern Kenya Rift (modified from Williams & Chapman 1986).

containing volcanic glass remains, are also intercalated in the red soil beds (Fig. 6c). Within the red soil beds there are also local gravelly sandstone beds (up to 1 m thick), which have a lenticular

shape and truncate the underlying red soil beds (Fig. 6b). Parallel stratifications without major bounding surfaces (Fig. 6b) and massive beds also occur in the sandstone beds.

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Fig. 2. Geological map of Samburu Hills (modified from Saneyoshi et al. 2006). (a) Regional geological map and east–west cross sections. Note that the Namurungule Formation has a scattered distribution within the study area. (b) Detailed geological map of the study site. NM2, NK5, KI1 and KI2 are the positions where the columnar cross sections in Figure 5 were measured. (c) and (c0 ) The detailed facies distribution within the Namurungule Formation and detailed section positions at which columnar cross sections were measured.

Red soil beds typically develop in dry climates with seasonal precipitation. The calcrete layers are caused by high evaporation of groundwater (cf. Retallak 1990; Khandkikar et al. 1999). The volcanic glass in the chert beds probably originated from volcanic ash. They are also interpreted as having formed in association with high evaporation (Tateishi 1987). The parallel stratified or massive sandstone beds without major erosion surfaces imply deposition from a single flood event, probably from an ephemeral stream. The massive sandstone beds are interpreted as having been deposited from low-density debris flows or hyperconcentrated grain flows [see fig. 5.1 in Miall (1996)].

The thick pile of red soil implies that the base level continuously rose throughout the deposition of the upper part of the formation. The thick basalt layers suggest that accommodation space in fragmented basins filled with a large volume of basalt lava in a short time period. The upper part of the formation, therefore, accumulated under underfilled conditions (see Carroll & Bohacs 1999 for details). But the very early phase of basin formation probably occurred under overfilled or balanced-filled conditions because the sediment supply was concentrated in small, juvenile basins (see fig. 21 of Withjack et al. 2002; Schlische & Olsen 1990).

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Fig. 3. Stratigraphy and chronology of the Samburu Hills (modified from Sawada et al. 1998). The ages of individual magnetostratigraphic unit in millions of years ago are from Cande & Kent (1995). M.U., Minor unconformity.

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Fig. 4. Columnar cross section of the upper part of the Aka Aiteputh and Namurungule formations, together with its magnetostratigraphy. The interpretation of depositional environments is based on Saneyoshi et al. (2006).

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Fig. 5. Columnar cross sections and results of the palaeoflow measurements and petrographic analysis of the Namurungule Formation. Q, quartz; M, microcline; F, K-feldspar; R, rock fragment. Abbreviations for columnar sections, C, channel fill deposit; F, delta front deposit.

Namurungule Formation The Namurungule Formation consists of alluvial fan and lacustrine delta sediments and has been already described in Saneyoshi et al. (2006) (Fig. 5). The

facies character is briefly reviewed below, except for the playa facies, which were not described in detail in Saneyoshi et al. (2006). The lowermost part of the Namurungule Formation is represented by conglomerate beds and interbedded sandstone

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Fig. 6. Outcrop pictures of the upper part of the Aka Aiteputh Formation. (a) Typical outcrop feature of the upper part of the Aka Aiteputh Formation, consisting of red soil beds (R) and frequently interbedded calcrete beds (CR). The outcrop is about 6 m high. (b) A close-up picture of the red soil beds and calcrete beds. Red soil beds rarely contain lenticular gravelly sandstone beds of ephemeral channel origin (CH). Calcrete beds show generally planar stratification but contain abundant root remains and show slightly irregular shapes in some locations, like this picture. (c) Chert beds (CT) interbedded here.

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beds of alluvial fan origin (Fig. 5). Clasts in the conglomerate beds are basalt fragments and debris from pyroclastic falls, such as pumice grains. The conglomerate facies is overlain by pyroclastic flow deposits (c. 5 m thick), which were dated to 9.6 Ma by Sawada et al. (1998), below the alternating sandstone and mudstone beds. Each sandstone and mudstone bed in the alternation is up to 0.5 m thick, and 0.05–0.1 m thick beds are most common in the lower part of the formation (Fig. 7a). The sandstone beds are characterized by parallel stratification, hummocky cross-stratification (HCS), and wave ripple lamination (Fig. 7a, b). The mudstone beds contain rootlets and burrows, and desiccation cracks are found on the top of some mudstone beds (Fig. 7c). Footprint traces, such as Carnivora, Artiodactyla and Rhinocerotidae, were also discovered on the sandstone surfaces (Nakano et al. 2001). Conglomerate beds (up to 0.5 m thick, averaging 0.2 m thick) showing sheet-like geometry occur rarely within the alternation. The conglomerate beds exhibit trough cross-stratification or parallel stratification with almost flat bases, and are topped in many cases by wave ripple lamination (Fig. 7d). There is no evidence of amalgamation of conglomerate beds in the lower part of the formation (see Fig. 7d). A sequence of sandstone with wave-generated sedimentary structures (wave ripple laminations and HCS) and mudstone records the subaerial exposure of the beds following their deposition in a subaqueous environment (Saneyoshi et al. 2006). The sandstone deposition was presumably influenced by wave or unidirectional flow. The overlying mudstone is also interpreted as having been deposited under calm subaqueous conditions once the wave or current activity ceased. The rootlets and desiccation cracks in the mudstone are suggestive of subaerial exposure following deposition. The successions are here interpreted as a flood plain deposit around an ephemeral lake, which was submerged temporarily by large-scale flooding (a larger expansion of the ephemeral lake). The HCS beds suggest inundation of the flood plain probably about 1–2 m. The HCS beds may have formed in a shallow-water environment as shown by the report of HCS beds from the modern lake surf zone (Lake Huron) shallower than 2 m (Greenwood & Sherman 1986). Conglomerate beds showing sheet-like geometry are interpreted as products of a single ephemeral stream (fluvial stream) deposit, which was then inundated following a lakelevel rise. In the southern part of the study area, greencoloured laminated mudstone, which is interpreted as a lacustrine (prodelta) deposit, is the dominant facies (Fig. 5). Each lamina is up to 1 mm thick. It is thickest in the south of the study area and

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tends to thin towards the north (Fig. 5). It appears around section NK5 in the sandstone and mudstone alternations, where the lower part of the Namurungule Formation becomes locally thicker than in the surrounding area (Fig. 5). The distribution of the lacustrine facies in the cross-section obviously shows a retrogradational stacking pattern (Fig. 5). The local appearance of the lacustrine facies around section NK5 is interpreted as a result of the tectonic subsidence being greater than in the surrounding areas. Saneyoshi et al. (2006) interpreted the interval above the conglomerate beds of the lower part of the Namurungule Formation as indicative of a delta system (Type 1 delta). The upper part of the Namurungule Formation is made up of sediment cycles characterized by laminated, green mudstone beds (prodelta deposit), a tabular cross-stratified sandstone bed (delta front deposit), and a trough cross-stratified sandstone bed showing a lateral accretion pattern (from a meandering stream deposit) and covered by rooted mudstone beds (from a flood plain deposit). Saneyoshi et al. (2006) interpreted this interval as indicative of a deltaic deposit, whose characteristics are different from the lower one (Type 2 delta). This difference was inferred because of the autogenic controls of the basin expansion (Saneyoshi et al. 2006). Playa deposit in the lower part of the Namurungule Formation. Just below the green laminated mudstone bed facies in sections KI1, KI2, and NK5, sediments consisting of fine alternations of a laminated green mudstone layer, a mud flake layer, and a laminated coarse siltstone layer (Fig. 8a, b) form a sediment interval to 1 m thick. The laminated green mudstone layer is about 1 –3 mm thick, and some of the beds are partly disturbed by desiccation cracks (Fig. 8c). The flakes in the mud flake layers (1– 10 mm thick) are platy in shape, and their approximate size ranges from 1–5 mm in diameter. Each laminated coarse siltstone layer is up to 20 mm thick. Rootlets were not found in the successions. The thickness of a single set of the layers mentioned above is 2–30 mm. The resemblance of the green mudstone beds to lacustrine laminated mudstone beds suggests they were deposited under subaqueous conditions, and the desiccation cracks indicate drying up of the lake at this site. The brecciated mud flakes seem to have been caused by desiccation crack formation, which might have caused the surface of the mud layer to peel prior to fragmentation and subsequent erosion by waves and currents when the lake re-expanded (desiccation breccia: Talbot & Allen 1996: see also Renaut et al. 1999 describing similar sediments as mud breccia from lake deposits of Tugen Hills, Kenya). The overlying laminated coarse siltstone layers might have been deposited

118 T. SAKAI ET AL. Fig. 7. Outcrop pictures of the lower part of the Namurungule Formation. (a) Alternation of sandstone and mudstone, which occupy the major part of the formation. In the lower part of the formation, hummocky cross-stratification is commonly seen in the sandstone beds. The mudstone generally contains rootlets and desiccation cracks (Fig. 7c). The outcrop is about 1 m high. (b) Wave ripple lamination in the sandstone beds. The scale is 1 cm in diameter. (c) Desiccation crack found on the top surface of a mudstone bed. The scale is 25 cm long. (d) Gravel beds of fluvial stream origin (G), alternating with silty sandstone beds. The arrow indicates the wave-ripple laminated sand covering the top of the gravel beds that suggests inundation of the channel. Although the wave ripple lamination was not preserved in the top of the beds, other gravel beds are also interpreted to be short-lived and submerged soon after their accumulation. The outcrop is about 1 m high.

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Fig. 8. Simplified columnar cross-section of the triplet composed of laminated mudstone beds, mud flake beds, and laminated coarse siltstone beds, which suggest repeated drying and inundation of the playa environment. (a) Simplified columnar cross-section of the beds. (b) The beds at outcrop. Arrows indicate the green-coloured laminated mudstone beds of lacustrine origin. White layers represent very fine to coarse siltstone beds with parallel lamination. MF, mud flake layers. (c) Desiccation crack developed in a green-coloured laminated mudstone bed.

from a weak stream or waves as the lake level rose. The deposit is, therefore, interpreted as being from a playa environment in which there was repeated inundation and drying out of the lake. This environment was probably spread throughout the interfluvial area of the lacustrine deltas. The beds do not contain any evidence of evaporite precipitation and dissolution, such as sand patch fabric. Namurungule Lake therefore appears to have been a fresh or nearly freshwater lake. The duration of deposition of a single set of layers mentioned above is difficult to estimate because the error range of the derived ages from the lower part of the Namurungule Formation is on the order of 0.1 Ma. A rough estimate of the average sedimentation rate is about 1.52 mm/ annum from the N2 interval in the Section NK5 (Saneyoshi et al. 2006). The set thickness is thicker than the annual average sedimentation rate,

but the sedimentation rate calculated here involves lacustrine sediments, accumulated under a slower sedimentation rate. Because the sedimentation rate at the lake margin appears to have been faster than at the lake floor, the sedimentation of a set a few millimetres thick could, therefore, be explained by annual lake-level rise and fall events driven by rainy (probably a long rainy season) and dry periods. The lake expansion attributed to the deposition of the alternations might have been caused by large-scale, lower-frequency flooding, whose exact causes are difficult to assess.

Provenance of the sandstones Sandstone petrography, together with visual observations of conglomerate beds and sandstone beds whose grains were coarser than medium sand, was used. Visual observation of the gravel and sand

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grains in the field showed that no grains from basement rocks (i.e. gneiss and granite) occurred in the sandstone beds of the upper part of the Aka Aiteputh and the lower part of the Namurungule Formations. Instead, the beds comprise basalt and pumice grains. In contrast, grains from basement rocks are abundant in the upper part of the Namurungule Formation. To confirm the absence of sand grains sourced from the basement rocks, a sandstone bed from the uppermost portion of the lower part of the Namurungule Formation was tested. Subsequently, the sandstone beds from the upper part were analysed to determine the interval at which grains from the basement rock first appeared. We identified minerals and the percentage of quartz, microcline, K-feldspar and rock fragments in 500 grains under the microscope. Quartz and microcline are strong indicators of a sediment supply from basement rocks, in contrast, K-feldspar and rock fragments were originated from the underlying Aka Aiteputh Formation or were supplied via pyroclastic falls. In total, ten samples were analysed from two different sections (KI2 and NK5) in the upper part of the formation. The results showed that specimens from the lower part consist only of K-feldspar and rock fragments. On the other hand, the quantity of quartz and microcline grains increases systematically up the stratigraphic column in the upper part of the formation in both sections (Fig. 5). This suggests that the catchment reached the basement area at that time when deposits of the upper part of the formation started to accumulate.

rock area (Fig. 2). However, the distance between the basins and the basement source rock area is thought to have been closer in earlier times, because the basins have continued to fragment as a result of the rift valley extension. The basin catchment at the time of deposition was very narrow, and the sediments and water reaching the basin were sourced from areas very close to the basins in the upper Aka Aiteputh and the lower Namurungule phases. A topographic high that strongly affected the precipitation is not thought to have been present around the Samburu Hills at c. 10 Ma for the following reasons: (1) Morley (2002) estimated the position of the doming centre at 10 Ma around the central part of the Kenya, about 50 km south of the Samburu Hills; (2) the large volcanic cone in the Aka Aiteputh Formation was absent; and (3) the rift basins in the Namurungule phase do not have the large topographic relief that is expected from the predominance of muddy facies on the delta plain environment, which suggests a smaller altitude difference in the past between the rift shoulder and basin floor. Because there is little possibility that the precipitation in the basin and the catchment area was subject to topographic control, the basin analysis here is suitable for discussing the regional climate prevailing in the northern part of the Kenya rift. The small catchment area of the basin also helps to detect the precise climate signals in the area, because the water supply was not affected by precipitation in areas far away from the basin. As a result, the climate in the basins was largely controlled by regional patterns such as the African Monsoon.

Discussion Palaeogeography of the basin and its significance for climate change reconstructions from the Samburu Hills Topographic differences are one of the most important controls on precipitation in present-day East Africa (e.g. fig. 2 in Bergner et al. 2003; Chan et al. 2008). Precipitation is larger around high mountains, and a drier climate spreads across lower altitude areas (lower than c. 1500 m a.s.l.). In East Africa, an area lacking high-altitude features around a basin is suitable for climatological analysis, so that the evolution of the climate system associated with the regional wind system can be understood. The results of the petrographic analysis reveal that the sediment and water supplies to the basin were only limited within the area between the basement rock area and the basins during the upper Aka Aiteputh and the lower Namurungule phases. Currently, in the Samburu Hills, the location of the target basin is about 15 km away from the basement

Climate shift from the Aka Aiteputh to Namurungule phase Before discussing the climate change from the Aka Aiteputh to the Namurungule Formations, differences in sedimentation rate should be considered, because soil formation is also a function of sedimentation rate, as well as environment. In the upper part of the Aka Aiteputh Formation, the calculated sedimentation rate is about 0.50 m/ka, based on the thickness divided by the total duration of sediment accumulation (Fig. 4). The rate for the lower part of the Namurungule Formation was estimated at about 1.52 m/ka, and for the upper part of the formation at about 0.24 m/ka around section NK5 (Saneyoshi et al. 2006). The sedimentation rate of the upper Aka Aiteputh phase may be attributed to rapid lava accumulation. Because about 50% of the lithofacies in the upper part of the Aka Aiteputh Formation are lava layers, the sedimentation rate is estimated to be smaller (at most 0.25 m/ka) if the lava thickness is removed from the calculation

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of sedimentation rate. This value is smaller than that of the lower Namurungule Formation, but is equal to that of the upper Namurungule Formation. The nearly similar sedimentation rates between the upper part of the Aka Aiteputh Formation and the upper part of the Namurungule Formation allow the lithofacies change between these formations to be interpreted as a result of a climate shift. Absence of lake deposits during the Aka Aiteputh period suggests that evapotranspiration was larger than water supply to the basin. As mentioned above, calcrete layers are generally associated with seasonal precipitation and subsequent high evaporation. The appearance of the lake and its enlargement during the Namurungule phase reveals that the precipitation was greater than evapotranspiration: the precipitation increased after the Aka Aiteputh phase. This result contradicts the vegetation reconstruction based on mammal fauna analysis in the Samburu Hills (Nakaya et al. 1987): open forest or savanna grassland is believed to have spread during the Namurungule phase. This reconstruction also proposed that the vegetation for the lower part of the Aka Aiteputh Formation was closer woodland. Although there are few mammal fossils from the upper part of the formation, the remains of larger rootlets imply the presence of woodlands or scattered woodlands in this phase. The difference in the reconstructed Namurungule phase climate change between the two methods is explained by the intensification of seasonality: heavier rainfall occurred during the rainy season, and drier conditions occurred during the

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dry season. Siliciclastic sediment input to basins may have been highest under conditions in which precipitation varies strongly with the seasons because they restricted vegetation cover in upland areas (e.g. Wilson 1973; Cecil 1990). In contrast, sediment supply is suppressed under dry conditions because of the absence of major streams and under wetter conditions such as rainforests, by deeper vegetation (Cecil 1990). The larger yield of clastic sediments in the catchment area caused by strong seasonality and pyroclastic falls might have suppressed forest development around the basin in the Namurungule phase.

Comparison with the surrounding area Sedimentary records dating to around 10 Ma in Kenya are very rare. The exceptional Ngorora Formation (13.1 –9.3 Ma, Deino et al. 1990; Hill 1999), exposed in the Tugen Hills (Fig. 1), has been well studied. It consists lithostratigraphically of five members (A– E). Members C and E are made of lacustrine facies (Hill 1999). The lake of the member C phase (around 12 Ma) was alkaline (Fig. 9), and its sediments record frequent drying of the lake water. In contrast, a nearly fresh lake water is inferred during the member E phase (11.5 –9.3 Ma; Fig. 9) (Hill 1999). Plant fossil analysis reconstructed a tropical rainforest around 12 Ma and an arid woodland around 10 Ma (Fig. 9) (Jacobs & Kabuya 1987; Kingston 1999). Kingston et al. (1994) carbon isotope analysis for pedogenic nodular materials revealed a C3-C4 mixed plant environment prevailed throughout the past 15 Ma, but

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Major doming

Tectonics & volcanic activity

Onset of rifting

Basin fragmentation

(Samburu Hills) Volcanism (basalt & trachyte)

(Maboko)

Climatic information

Dry open woodland

Volcanism (basalt & minor trachyte) Tilting of the Namurungule basins

Climatic shift

Alkaline lake

Fresh water lake

(Tugen) Rainforest

Arid-open woodland

Seasonal woodland - Rainforest

(Samburu Hills) Forest/closer woodland

Savanna/open woodland + lake

Fig. 9. Tectonic and climate events in central to northern Kenya in the past 20 Ma. Complied based on Jacobs & Kabuya (1987), Nakaya et al. (1987), Sawada et al. (1998), Hill (1999), Kingston (1999), Renaut et al. (1999), Retallak et al. (2002), Pickford et al. (2004), Sawada et al. (2006).

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analysis of mammal tooth enamel apatite delineated a shift from C3- to C4-dominant plant phases between 8–6 Ma, which correlates to the worldwide C3-to-C4-shift (cf. Cerling et al. 1989). However, Kingston never detected a climate shift around 10 Ma in the flora. Recently, Kunimatsu et al. (2007) reported finding a new hominoid fossil from the fluviolacustrine sediments (c. 9.9 Ma) at Nakali, located between the Tugen and Samburu Hills. This period is in the Aka Aiteputh phase, suggestive of a wetter climate than is currently found in the Samburu Hills. From the data sets from both Tugen Hills and Nakali, we conclude that a wetter climate was found south of the Samburu Hills before 9.6 Ma (Aka Aiteputh phase). Local climate differences between the Samburu Hills and the central part of Kenya (the Tugen Hills and Nakali) might be due to palaeogeographic differences associated with the regional rock upliftdoming processes that occurred after 15 Ma (cf. Chorowicz 2005). The latter two areas were located near the doming centre in Kenya at that time (cf. Morley 2002), where the palaeoaltitude is expected to have been higher than in the Samburu Hills. These two areas seem to have had greater precipitation due to higher altitude as in present-day Kenya (e.g. Bergner et al. 2003): the palaeotopographic difference resulting in the difference of climate between the Samburu Hills and further south (Tugen Hills and Nakali). Although distinct evidence for climate change has not been detected from isotopic analysis of materials from 10 Ma in the Tugen Hills, the lake water change from alkaline to fresher water may record increased precipitation in the member E phase (after around 11.5 Ma), slightly earlier than in the Samburu Hills. The shift from tropical rainforest to arid woodland (cf. Kingston 1999) seems to have occurred at almost the same time as the shift in the Samburu Hills, suggestive of shift to dry climate. The data set implies the possibility of the intensification of seasonality also happened to the south of Samburu Hills around (or slightly before) 10 Ma.

Probable causes of the climate shift in 10 Ma As discussed below, the climate shift detected in the Samburu Hills occurred at the same time as that in the Himalayas and the surrounding area. This phenomenon may be associated with the intensification of the Indian summer monsoon (ISM), which has been recognized as the manifestation of the Intertropical Convergence Zone (ITCZ) during boreal summer (e.g. Gadgil 2003). Before discussing the probable causes for the climate shift in the

Samburu Hills, we briefly summarize the present climate of Kenya. The present dry climate prevailing in most of the Kenya rift is associated with the area’s uplift since 20 Ma (e.g. Sepulchre et al. 2006; Spiegel et al. 2007) as well as the global effect such as Indian Ocean surface temperature cooling and the onset of the glacial and interglacial cycles (e.g. Cane & Molnar 2001; deMenocal 2004). The precipitation pattern is similar to that in other tropical regions, which experience two rainy seasons. Kenya is, therefore, categorized in the equatorial-tropical precipitation zone (Gasse 2000), although more severe dry conditions continue at lower altitudes during dry seasons. Precipitation occurs in a spring season (long rainy season) and autumn (short rainy season) when the ITCZ passes from the south to the north and vice versa. The moisture is transported both from the Atlantic and Indian Oceans in the rainy seasons (Chan et al. 2008). The divergent ISM wind in the summer season, in contrast, brings dry air to most of the northern Kenya Rift (e.g. Camberlin 1996). The topography also controls precipitation within the rift valley region, as mentioned above, and higher areas such as Mt Kenya and the Aberdare Range are subject to heavier precipitation (e.g. Bergner et al. 2003), even during the dry season (e.g. Vincent et al. 1979; Camberlin 1996). Present-day Kenyan precipitation is mainly controlled by ITCZ movement, and appears to have no direct relationship with the ISM. However, the ITCZ strength over Kenya in the spring season and the subsequent ISM wind is related. Strong ITCZ rainfall in East Africa in the long rainy season is followed by heavy rainfall in Himalayan region from a strong ISM. Weaker ITCZ activity is followed by a weaker ISM, coupled with systematic sea surface temperature change in the Indian Ocean (e.g. Webster 2006). Similarly, a strong positive correlation of summer monsoon wind strength and ITCZ activity is also revealed in the Asian region (e.g. Yancheva et al. 2007). In late Pleistocene records, wet–dry cycles in East Africa are basically synchronized with ISM activity. The precipitation maximum was estimated from the lake-level record of Lake Naivasha in central Kenya to be 9–11 and around 135 ka (Trauth et al. 2001; Bergner et al. 2003; Trauth et al. 2003). In particular, the 135 ka highstand event was also reported from other major Kenyan lakes (Trauth et al. 2001). These studies interpreted that the 9–11 ka highstand is correlatable to the June –July insolation maximum in the Northern Hemisphere (Bergner et al. 2003; Trauth et al. 2003). In contrast, the maximum March insolation on the equator was inferred as the cause for the 135 ka highstand (Bergner et al. 2003; Trauth et al. 2003). Around the Indian Ocean, the rapid

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deposition of speleothem with lower O isotope ratio occurred in the period 120 –135 ka in northern Oman, suggestive of an increase in precipitation due to the northward shift of the monsoon precipitation belt (Fleitmann et al. 2003), that is, intensification of ISM activity. This period roughly corresponds with the lake-level highstand period in the East Africa. Lake-level records from other East African lakes also indicate rough correspondence with the global glacial –interglacial cycle: lake-level lows occurred during the glacial periods and highs during interglacial periods (Gasse et al. 1989; Cohen et al. 1997; Stager et al. 2002). These lake-level records and other climate proxies imply a dry climate prevailed during glacial periods and a wetter one dominated during interglacials. The intensification of northern hemisphere summer insolation is thought to be attributable to enhanced ITCZ convection, as well as the African and Indian summer monsoons (cf. Le´zine & Casanova 1991; Barker et al. 2002; Bergner et al. 2003; Le´zine et al. 2005), suggestive of the close relationship between the ITCZ and ISM in the past. Before 2.7 Ma when a dry climate system was established in East Africa (Cane & Molnar 2001; Trauth et al. 2005; Bobe 2006), climate records are sparse and are insufficient for discussing ITCZ and IMS control on the East African climate. Aridification accompanying grassland spreading already started before 2.7 Ma and extends, at least, back to around 8 Ma (Cerling 1992; Se´galen et al. 2007), related to East African topographic uplift that blocked moisture transport both from the west and east (e.g. Sepulchre et al. 2006), as well as the closure of Indonesian seaway around 3–4 Ma (Cane & Molnar 2001). Owing to the lack of climate record and palaeogeographic information in East Africa around 10 Ma, it is difficult to discuss the exact causes of the climate shift at that time. But the timing of the shift is almost simultaneous with the onset of the intensification of the ISM that was detected in the western Indian Ocean, offshore Oman, where strong upwelling occurs because of the monsoon wind (e.g. Gupta et al. 2003), as well as at the foot of the Himalayas (Siwaliks; Barry et al. 1995; Nakayama & Ulak 1999; Dettman et al. 2001; Ulak & Nakayama 2001), in the Tibetan Plateau (e.g. Dettman et al. 2003), and in the Indian Ocean and surrounding basins (Rahman & Ruth 1990; Kroon et al. 1991). Chemical and palaeoceanographic data implies that the intensification started around 10.7 Ma (Dettman et al. 2001; Rahman & Ruth 1990) and fluvial facies changed slightly after that time (Nakayama & Ulak 1999). The timing of facies changes in the Samburu Hills is consistent with changes observed in the Siwalik Hills.

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One reason that complicates discussing the causes of the climate shift is our poor understanding of palaeogeography around 10 Ma. Northern Kenya is thought to have been located slightly further south around 10 Ma than it is now (e.g. George et al. 1998); the wind system that controlled the climate of the region at that time is thought to have been similar to the present system. Recent meteorological model experiments by Sepulchre et al. (2006) indicated the possibility of an increase in precipitation driven directly by IMS intensification in East Africa before any topographic uplift. In their experiments, the easterly trade wind during boreal summer could have penetrated deep into the African continent, transporting much more moisture to East Africa than is the case now, supposing the Ethiopian Rift altitude (present altitude c. 3000 m a.s.l.) was lower than it is now (cf. Pik et al. 2008). This would have led to rainfall during the summer Indian monsoon season (July– August). As for our results, there are two possible scenarios explaining the climate shift detected in the Samburu Hills: (1) both the ITCZ activity and ISM, linked as in the present climate system, may have intensified; and (2) the intensified ISM over the East African region during boreal summer caused by lower altitude could have allowed moisture penetration from the ISM deep into East Africa. It is impossible to determine which scenario most likely drove the climate shift detected in the Samburu Hills, because we cannot know whether the palaeoaltitude at that time was high enough to block moisture transport from the east and west to inland East Africa. But scenario (1) seems to be plausible in explaining the climate shift, because scenario (2) would drive a wet annual climate, as in other tropical regions, and it cannot explain the increase of seasonality. As mentioned above, the wet climate indicated in the Tugen Hills and Nakali was probably due to their higher altitudes, which resulted in the air stream encountering topographic obstacles and causing greater precipitation, as in modern high altitude areas. However, tropical forests have been proposed in the Tugen Hills, just before the climate shift happened (Jacobs & Kabuya 1987). The presence of tropical forests at that time implies that the central Kenya Rift had a lower altitude then, allowing the stronger moisture transport from the oceans to the area. In contrast, Pickford (1987) reported gastropod fossils from the Namurungule Formation whose present habitat is 1400– 1600 m a.s.l., and he inferred the similar altitude for the Namurungule basins. Retallak et al. (2002) estimated palaeoprecipitation of 300–500 mm for 15 –13 Ma based on pedological analysis around Maboko Island on the

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east coast of Lake Victoria. They estimated heterogenic forest, including grassland, in the area at that time. Retallak (1991) showed evidence of grassy woodland or woody grassland from the c. 14 Ma succession at Fort Ternan, southwestern Kenya, suggestive of dry climate. These may suggest uplift of the area at that time (cf. Retallak 1991). The altitude at that time was probably lower (cf. Pickford 1990), but it might have been high enough to suppress the penetration of the easterly (ISM) into East Africa in boreal summer until around 10 Ma. The newly discovered record of climate shift from the Samburu Hills is the first record that is almost synchronous with the early phase of ISM intensification, but the complicated climate system (experiencing effects of both Indian and Atlantic Oceans) together with a poor understanding of local palaeogeography makes discussing the exact mechanism of climate change difficult. More information about the climate is needed from the surrounding area for a more reliable analysis of the regional climate system and geography around 10 Ma.

Conclusions A climate shift was detected from the Middle to Upper Miocene Aka Aiteputh and Namurungule Formations, in which a series of filled small half grabens developed in the early phase of the development of the northern Kenya Rift. The estimated climate during the upper Aka Aiteputh phase (c. 10.1–9.6 Ma) was dry with seasonal precipitation, as suggested by abundant calcrete layers in red soil beds. The Namurungule phase climate was, in contrast, characterized by strong seasonality. In particular, precipitation was greater then than during the Aka Aiteputh phase. Petrographical study results, together with the absence of intense volcanic activity nearby, suggest the catchment area of the basin was limited to the zone around the basin. There seem to have been no high mountains that could have affected regional climate through topographic effects, and the climate record of the basin reflects that of the regional climate. The climate shift that occurred around 9.6 Ma is similar in age to that detected around the Himalayas, Tibet, and the Indian Ocean, in particular, offshore Oman. The probable causes of the shift are (1) intensification of ITCZ convection synchronized with Indian summer monsoon intensification, and (2) deep penetration of Indian summer monsoon wind into the rift valley if the altitude of the rifted area was lower than it is now. The former scenario seems to be the more plausible. Further climatological studies are required to understand the atmospheric circulation system of Africa in the Miocene.

The authors thank Prof. P. Clift (University of Aberdeen, UK) and Prof. R. Tada (University of Tokyo, Japan) who gave us an opportunity to publish our paper, and the Turkana people in Nachola Village who assisted our fieldwork in the Samburu Hills. The government of Kenya is appreciated for its kind permission for our research in Samburu Hills. JSPS Nairobi Office supported our field work in Kenya. Dr H. Nakaya (Kagoshima University, Japan) and Mr K. Uno (University of Utah, USA) are thanked for their discussions and encouragement. We thank two anonymous reviewers and the editor for our manuscript, Prof. H. Zheng (Nanjing University, China) for their critical comments. This work was supported by the grant-in-aid from Ministry of Education, Japan (17740335 for TS, 14253006 for YS and 19207019 for MN).

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Estimation of atmospheric CO2 uptake by silicate weathering in the Himalayas and the Tibetan Plateau: a review of existing fluvial geochemical data YOUNGSOOK HUH School of Earth and Environmental Sciences, Seoul National University, 599 Gwanak-ro, Gwanak-gu, Seoul 151-747, Korea (e-mail: [email protected]) Abstract: The fluvial geochemical data of major rivers draining the Himalayas and the Tibetan Plateau (HTP) are compiled from literature and supplemented with data from the author’s group to explore the spatial variability in the major element and strontium isotopic compositions and in rates of silicate weathering and concomitant drawdown of atmospheric CO2. The results indicate that carbonate weathering dominates the major element composition of these rivers, and when the silicate contributions are carefully separated, there is spatial variability within the HTP system. Silicate contributions are highest in the rivers of the Himalayan syntaxes and in the Yamuna, Alaknanda-Bhaghirathi and Kosi tributaries of the Ganges. The 87Sr/86Sr ratios are especially radiogenic in the Ganges tributaries but the values do not necessarily correlate with the relative input from silicate weathering. Even in the Yamuna tributary of the Ganges where rates of CO2 consumption by silicate weathering are several times those of the rivers of the eastern Tibetan Plateau, the rates are comparable to the rivers draining the Andes. Thus, the Ganges tributaries supply uniquely radiogenic 87Sr to the ocean but cannot be considered anomalous in terms of silicate weathering rates in comparison to other major rivers draining orogenic zones. Supplementary material: Major element concentration and strontium isotope ratio data for Huh (unpublished) is available at http://www.geolsoc.org.uk/SUP18408.

Uplift of the Tibetan Plateau, inclusive of the Himalayas and the associated mountain ranges of Central Asia, has occurred continuously over the last c. 40 Ma with a seemingly profound effect on the evolution of global climate during that time period. Regional elevations thus attained are currently the largest on Earth, and this has had large-scale impacts on atmospheric and oceanic circulation (Ruddiman et al. 1997). The active tectonics have generated steep slopes and focused precipitation on their margins, increasing physical and chemical weathering. Chemical weathering of silicate minerals consumes atmospheric CO2, and this uplift-weathering control of CO2 forms the crux of the hypothesis to explain the Cenozoic global cooling (Raymo & Ruddiman 1992). However, the nature and timing of the influence of the Himalayan orogeny on weathering and global climate are still contentious. The underlying assumption is that atmospheric CO2 is the main determinant for climate change over million year timescales. The major fluxes that affect atmospheric CO2 levels in the global carbon cycle are multiple. Mid-ocean ridge volcanism and metamorphic decarbonation in orogenic zones add CO2 to the atmosphere. During the Cenozoic era, the CO2 input flux from mid-ocean ridge volcanism probably remained steady. The palaeorecord of foraminiferal Li/Ca ratio indicates that

hydrothermal exchange has not varied by more than 30–40% (Delaney & Boyle 1986), and seafloor production rates were relatively constant (Rowley 2002). The imbalance between weathering and burial of organic carbon is another source of CO2. The major net sink is consumption of atmospheric CO2 during weathering of silicate [reaction (1) forward] or of carbonate [reaction (2) forward] minerals: CaAl2 Si2 O8 þ 2CO2 þ 3H2 O O Al2 Si2 O5 (OH)4 þ Ca2þ þ 2HCO 3 CaCO3 þ CO2 þ H2 O O Ca

(1) 2þ

þ

2HCO 3

(2)

The plagioclase (CaAl2Si2O8) of reaction (1) is only one example of Ca- or Mg- containing silicate minerals. Of the two units of atmospheric CO2 consumed during weathering of silicate minerals one is precipitated as CaCO3 after transportation by rivers as HCO2 3 to the ocean, and the other one unit is released back into the atmosphere as CO2 [reverse of reaction (2)]. Thus, weathering of Caor Mg-silicate minerals is a net sink for atmospheric CO2. On the other hand, weathering of CaCO3 consumes one unit of CO2 [reaction (2) forward], and the same amount is returned to the atmosphere upon precipitation of CaCO3 in the ocean [reaction

From: Clift, P. D., Tada, R. & Zheng, H. (eds) Monsoon Evolution and Tectonics –Climate Linkage in Asia. Geological Society, London, Special Publications, 342, 129–151. DOI: 10.1144/SP342.10 0305-8719/10/$15.00 # The Geological Society of London 2010.

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(2) backward]. Na- and K- containing silicate minerals pose some uncertainty. Their weathering reactions consume CO2 [e.g. reaction (3) forward]. It is commonly suggested that the Naþ and Kþ ions are transported by rivers to the ocean where they undergo reverse weathering to form authigenic clays [reaction (3) backward], in which case there is no net consumption of atmospheric CO2. 2KAlSi3 O8 þ 2CO2 þ 11H2 O O Al2 Si2 O5 (OH)4 þ 2Kþ þ 2HCO 3 þ 4H4 SiO4

(3)

Na- and K- containing clays are thermodynamically stable in seawater (Helgeson & Mackenzie 1970) and incubation experiments have shown that their authigenesis is possible (Michalopoulos & Aller 1995). However, marine sediments with significant amounts of authigenic clays have not been found except in near-shore environments. More likely, some alkali elements undergo ion exchange for alkali earth elements [e.g. reaction (4)] in which case there will be a net consumption of CO2 upon deposition of CaCO3 [reaction (2) backward]. Considered together, the Na- and K-containing silicate minerals are less efficient than Ca- and Mg-silicate minerals in CO2 consumption. 2Naþ þ Ca-clay O Na2 -clay þ Ca2þ

(4)

In the literature, the rate of CO2 consumption (fCO2) by weathering is expressed in diverse ways. On timescales shorter than precipitation of CaCO3 or reverse weathering, the CO2 taken up during weathering of carbonate and silicate rocks is: fCO2 carb ¼ f(Mgcarb þ Cacarb ) fCO2 sil ¼ f(Nasil þ Ksil þ 2Mgsil þ 2Casil )

(5) (6)

The subscripts ‘carb’ and ‘sil’ denote ions originating from carbonate and silicate minerals, respectively, and f indicates flux (mol a21) or yield (mol km22 a21). On timescales longer than about 1 million years, one has to consider the return of CO2 by CaCO3 deposition and reverse weathering, and assuming that reverse weathering is quantitative, the net consumption of CO2 is: fCO2 sil ¼ f(Mgsil þ Casil )

(7)

In this paper, I aim to address the CO2 sink potential of the Himalayas and the Tibetan Plateau (HTP) as revealed in the dissolved load of the rivers draining this area. To this end, the distinction and separation of carbonate from silicate weathering is of consequence. The results will contribute to global carbon cycle modelling efforts where the chemical composition of the major rivers is one of the main uncertainties (Kashiwagi et al. 2008).

To account for the variations in atmospheric CO2, the variability in sources and sinks other than weathering needs to be considered. For example, in some systems like the Himalayan front in Nepal, the input through metamorphic decarbonation reactions may be significant enough to negate the output by silicate weathering (Becker et al. 2008; Evans et al. 2008). The spatial and temporal variation of this decarbonation flux requires further consideration. Another important flux term concerns the organic carbon sub-cycle. The magnitude of CO2 drawdown by organic carbon burial surpassing that by silicate weathering was initially suggested from analyses of Neogene sediments of the Bengal Fan (France-Lanord & Derry 1997), where high organic carbon burial rates were sustained by the high sedimentation rate and low oxygen availability (Galy et al. 2007). In a comprehensive study of the Yamuna basin including the redox-sensitive element Re, the CO2 release from oxidative weathering of organic-rich sediments was found to be comparable to the CO2 draw-down by silicate weathering (Dalai et al. 2002a, b). Recognizing that there are other components in the global carbon cycle that may strongly influence the Cenozoic climate change and the role played by the Himalayas and the Tibetan Plateau, in this paper I focus on the inorganic carbon cycle involving weathering of silicate minerals. Uplift of the Himalayas and the concomitant rise in the marine 87Sr/86Sr record led many workers to make the connection between the two for the past 40 Ma (Richter et al. 1992). Although it is now widely accepted that Sr isotopes are not proxies of ‘general’ continental weathering (Edmond 1992; Huh & Edmond 1998; McCauley & DePaolo 1997), it could still function as an index of the weathering of ‘Himalaya-type’ orogenic belts. The radiogenic 87Sr/86Sr ratios and high Sr concentrations of the Ganges-Brahmaputra system (Krishnaswami et al. 1992; Palmer & Edmond 1992) seem to be unique to the distinctive kind of metamorphism associated with the continent—continent type collision. Determining how weathering in the HTP influences the marine record has important ramifications for the interpretation of the now beautifully constrained Phanerozoic marine 87Sr/86Sr record (Burke et al. 1982).

The scope of this study Objectives The objectives of the present study are to attend to the following questions. (1) Are silicate weathering rates and CO2 consumption rates accelerated in the HTP compared

CO2 UPTAKE IN THE HIMALAYAS AND THE TIBETAN PLATEAU

to other orogenic regions? What is the spatial variability within the HTP? (2) How general, spatially, are the enhanced radiogenic Sr fluxes observed in preliminary sampling of the Himalayan streams (Krishnaswami et al. 1992; Palmer & Edmond 1992)? Is there a relationship between fluvial Sr isotope ratios and silicate weathering or CO2 uptake rates?

Data sources The issue of weathering rates can be addressed from many angles: material balance in soil profiles, cosmogenic isotope measurement of erosion, field data from small streams, laboratory dissolution experiments, and so on. The analyses presented in this paper are based on dissolved load data sets for rivers draining the four sides of the HTP (Fig. 1, Table 1). They are medium size tributaries feeding

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into the major rivers of the area –the Salween, Mekong, Chang Jiang (Yangtze), Hong (Red), Huang He (Yellow), the internally drained rivers of the Tibetan Plateau interior and Tien Shan, Indus, Ganges, and Brahmaputra (Fig. 1). These have the advantage of averaging over a relatively large area and can be applied directly to current silicate weathering rates without scaling up. They represent ‘chemical’ weathering rates, which is the mechanism by which atmospheric CO2 is consumed and which cannot be estimated from physical weathering rates (Summerfield & Hulton 1994). The disadvantages, on the other hand, are the inevitable intra-basin variations in lithology and environmental parameters. The uncertainties arising from this variability are discussed further in the section ‘uncertainties’. Since the pioneering work by the group at the Physical Research Laboratory of India (Sarin &

Fig. 1. Location map showing the Himalayas and the Tibetan Plateau and the major rivers that drain it. Area with elevation greater than 2 km is shaded. The stars mark the furthest downstream samples included in this study. Only data from above these locations were used in order to specifically account for weathering in the mountains v. processes in the plains. Readers are referred to the original papers in the literature for exact sample locations of individual samples (Table 1).

132

Table 1. Description of data sources for each river system of the Himalayas and the Tibetan Plateau River system Salween Mekong Chang Jiang Hong (Red) Huang He Internal

Ganga –Yamuna Ganga –Bhagirathi–Alaknanda Ganga –Ghaghara

spot samples (23 summer, 1 winter) spot samples (19 summer, 11 winter) spot samples (55 summer, 0 winter) spot samples (23 summer, 20 winter) spot samples (20 summer) spot samples (38 summer) spot samples (125 summer, spring, fall; 33 winter) spot samples (22 summer, 7 monsoon, 26 post-monsoon) spot samples (134 non-winter but multiyear) spot samples

Discharge data type

Silicate calculation

individual samples

(Ca/Na)sil ¼ 0.35 (Mg/Na)sil ¼ 0.24 (Ca/Na)sil ¼ 0.35 (Mg/Na)sil ¼ 0.24 (Ca/Na)sil ¼ 0.35 (Mg/Na)sil ¼ 0.24 inverse model

Noh et al. 2009 Ellis et al. unpublished data Moon et al. 2007

individual samples

inverse model

Wu et al. 2005

(Ca/Na)sil ¼ 0.35 (Mg/Na)sil ¼ 0.24 Casil þ Mgsil ¼ Cariver þ Mgriver 20.5*HCO3 river (Ca/Na)sil ¼ 0.7 (Mg/Na)sil ¼ 0.3

Huh, unpublished data*

(Ca/Na)sil ¼ 0.7 (Mg/Na)sil ¼ 0.3 (Ca/Na)sil ¼ 0.2

Krishnaswami et al. 1999 Bickle et al. 2003 Galy & France-Lanord 1999 Galy et al. 1999 English et al. 2000 France-Lanord et al. 2003 Galy & France-Lanord 1999 Galy et al. 1999 Oliver et al. 2003 Quade et al. 2003 Hren et al. 2007 Huh, unpublished data*

individual samples individual samples individual samples

– this work select samples select samples select samples

(Mg/K)sil ¼ 0.5 Ganga –Gandak

spot samples (49 representative of whole year)

select samples

(Ca/Na)sil ¼ 0.2 (Mg/K)sil ¼ 0.5

Gaga –Kosi

spot samples (93)

this work

Brahmaputra

spot samples (176 representative of whole year)

individual samples

(Ca/Na)sil ¼ 0.41 (Mg/Na)sil ¼ 0.24 (Ca/Na)sil ¼ 0.545 (Mg/Na)sil ¼ 0.30

*Data are presented as Supplementary Publication.

Data source Noh et al. 2009 Noh et al. 2009

Karim & Veizer 2000 Pande et al. 1994 Huh, unpublished data* Dalai et al. 2002a Dalai et al. 2003

Y. HUH

Indus

Sampling interval (Number of samples)

CO2 UPTAKE IN THE HIMALAYAS AND THE TIBETAN PLATEAU

Krishnaswami 1984), systematic data have been collected on the fluvial geochemistry of large unperturbed river basins in the HTP, the main-stems as well as their many tributaries. This has allowed insight into the rates of weathering fixation of CO2 as a function of climate, lithology, topography and other potentially important environmental variables. Coverage is most comprehensive for the Ganges in India and Nepal, augmented recently for the Brahmaputra, the three rivers of eastern Tibet (Salween, Mekong, and Chang Jiang), the Hong (Red River) in southwestern China and Vietnam, and Huang He in northwestern China. To build a comprehensive data set (.900 samples), data on the headwaters of the Indus and the interior of the Tibetan Plateau (Huh, unpublished data) are added, and these are available as Supplementary Material. The published data sets that I include here (Table 1) for the Salween, Mekong and Chang Jiang (Noh et al. 2009), Hong (Moon et al. 2007), Huang He (Wu et al. 2005), Indus (Karim & Veizer 2000; Pande et al. 1994), the tributaries of the Ganges: Yamuna (Dalai et al. 2002a, 2003), Alaknanda-Bhagirathi (Bickle et al. 2005; Krishnaswami et al. 1999), Ghaghara (English et al. 2000; Galy & France-Lanord 1999; Galy et al. 1999), Gandak (France-Lanord et al. 2003), and Kosi (Quade et al. 2003); and the Brahmaputra (Hren et al. 2007; Singh et al. 2005, 2006) are by no means exhaustive. Only data covering the HTP proper or in the very foothills were selected and not the major tributaries that have significant inputs from the Siwaliks or the plains (Fig. 1). I have also excluded hot and cold springs and very small streams (drainage area of less than tens of km2), though they may be of interest for other purposes, and have based my estimates on river systems sampled at regional scales. Data sets are also limited to those with a complete set of major elements including silica, strontium, and 87 Sr/86Sr. Regarding the samples from the author’s own lab, an extensive suite of samples has been collected from the headwaters of the Indus in northern Pakistan, from the Indus and Brahmaputra tributaries in Tibet, headwaters of the Ganges tributaries in Nepal, and in western China by colleagues during geological and geophysical expeditions. These are ‘grab’ samples collected in pre-cleaned polyethylene bottles and filtred in the laboratory months later. As extensive storage tests comparing unfiltred Amazon samples showed, in sedimentpoor rivers untreated samples maintain their sample integrity as regards the elements considered here, and in sediment-rich rivers there may be a 5% increase in untreated samples (Stallard 1980).

133

Description of the study area Only a brief overview will be provided here and readers are referred to the original papers for more detailed information. The Tibetan Plateau has a mean elevation of c. 5 km over an area of c. 4  106 km2. The southern and northern boundaries have steep topographic fronts, but the southeastern and northeastern parts have continuous slopes (Burchfiel 2004). There is a sharp climatic contrast between the Tibetan Plateau and the High Himalayan range. The former experiences cold and arid climate with precipitation around 200– 500 mm a21; the latter is exposed to monsoons bringing precipitation 1000– 4000 mm a21 (Leemans & Cramer 1991). Major Asian rivers have sources in the HTP and flow around the two syntaxes, one in the Nanga Parbat and the other in the Namche Barwa (Fig. 1). The northern part is arid and the streams there are intermittent and internally drained. The rivers of the eastern Tibetan Plateau, the Salween, Mekong, and Chang Jiang, drain mainly the flysch sediments with scattered exposures of plutonic and basaltic rocks along the suture zones (Noh et al. 2009). The Hong drains the metamorphic rocks of the Ailao Shan-Red River Fault zone (Moon et al. 2007). The Huang He basin is composed of sandstone, limestone and evaporitebearing red beds (Wu et al. 2005). The internal drainage occurs in the Tien Shan Mountain and the Tarim Basin. Many of the Indus headwater streams drain exclusively igneous and metamorphic rocks of the Crystalline Core Complex or the ophiolites and trans-Himalayan plutons associated with the pre-collision active margin of southern Eurasia (Pande et al. 1994; Karim & Veizer 2000). The Himalayan front has similar structure longitudinally, with the Tibetan Sedimentary Series (TSS), High Himalaya Crystalline (HHC), and Lesser Himalaya (LH) from north to south, and the main tributaries of the Ganges flow across these units (Krishnaswami et al. 1999). The TSS is the former reef carbonate rocks of Asia that have undergone metamorphism and are composed primarily of Palaeozoic and Mesozoic carbonate and siliciclastic sediments. The HHC forms the high ranges and principally consists of ortho- and para-gneisses, migmatites and highly metamorphosed marbles. The Lesser Himalaya is composed of variably metamorphosed Precambrian sediments with quartzo-pelitic schists, quartzites and dolomitic carbonate rocks. The Brahmaputra flows along the Indus-Tsangpo suture zone and around the eastern syntaxis of Himalaya, where exhumation rates are up to 10 mm a21 (Zeitler et al. 1993; Burg et al. 1998). Thus, suites of samples are available for all the important geological units of the collision zone and the plateau.

134

Y. HUH

Calculation method for CO2 consumption rates

p , 0.001) (Moon et al. 2007). The two methods are briefly outlined below.

The dissolved load of streams has origins from atmospheric input, evaporite dissolution, and carbonate and silicate weathering. Quantifying exclusively the silicate contribution is critical, as that is the only net sink of atmospheric CO2 over timescales longer than a million years. In the published literature, each author uses different methods to calculate this, and their original calculations have been adopted here instead of reprocessing the data using a single canned method for all systems. This is because whichever method is used, detailed knowledge and educated assumptions are required of the regional environment. Each basin is distinct both lithologically and environmentally, and I believe that my efforts could not surpass those of the original authors who have carried out the fieldwork and carefully considered the characteristics of the basins. To present a consistent data set, the rate of CO2 consumption per unit area (CO2 yield, fCO2 in mol km22 a21) was recalculated as follows:

The forward model

fCO2 sil short-term ¼ (Nasil þ Ksil þ 2Casil þ 2Mgsil )  mean annual runoff  basin area (8) fCO2 sil long-term ¼ (Casil þ Mgsil )  mean annual runoff  basin area

(9)

If the timescales considered are shorter than those involved in carbonate precipitation and reverse weathering, fCO2 short-term would be a valid representation of the uptake rate of CO2 by silicate weathering. On timescales longer than the time it takes to return atmospheric CO2 consumed during weathering by carbonate precipitation in the ocean or reverse weathering, fCO2 long-term applies. Since there is some uncertainty regarding the effectiveness of Na- and K- silicate minerals, both approaches will be used (Table 2). Two methods have been used to calculate the silicate fractions: the forward and inverse models. In the literature, the Salween, Mekong, Chang Jiang, Indus, Ganges, and Brahmaputra data have been processed using variations of the forward model and the Hong and the Huang He data using the inverse model (Table 1). In the case of the Hong, both forward and inverse models were applied, and there was less than 9% difference in the mole percentages of SCatsil/SCattot ¼ (Na þ K þ Mg þ Ca)sil/(Na þ K þ Mg þ Ca)river and the two results were well-correlated (r2 ¼ 0.756,

The forward model sequentially subtracts components with pre-assigned compositions from the total riverine load for each element. There are variations of this model, but the most general form is outlined below. The atmospheric component is subtracted using the lowest Cl concentration in a system and known X/Cl (X ¼ Na, K, Mg, Ca) ratios of rain in the drainage basin. Sometimes this step is skipped altogether because the correction is minimal. The evaporite fraction is assumed to be mainly halite and gypsum, and Na and Ca are subtracted in stoichiometric proportion to the rain-corrected Cl and SO4 values. The derivation of the silicate fraction, which is key to calculating uptake of atmospheric CO2, is as follows (e.g. Noh et al. 2009): Nasil ¼ Nariver  Naatm  Naevap Ksil ¼ Kriver Casil ¼ (Ca=Na)sil  Nasil

(10) (11) (12)

Mgsil ¼ (Mg=Na)sil  Nasil

(13)

The values of (Ca/Na)sil and (Mg/Na)sil vary with local rock composition and therefore can differ by basin (Table 1). For large drainage basins with heterogeneous lithology, global estimates of Gaillardet et al. (1999) can be used: (Ca/Na)sil ¼ 0.35 + 0.15 and (Mg/Na)sil ¼ 0.24 + 0.12, based on modal values of silicate-draining small streams of North America and Europe (White & Blum 1995), the tropical rivers of the Guayana Shield (Edmond et al. 1995), and small streams in France (Meybeck 1986). The remaining cations after silicate correction are attributed to carbonate weathering.

The inverse model The inverse model assumes that input from four sources – rain, evaporite, carbonate, and silicate – can explain the dissolved major element and Sr isotope data (Negrel et al. 1993). The model consists of a set of mass balance equations: 

 X X  X ¼ ai,Na Na river Na i i

(14)

87    X 87 Sr   Sr  Sr Sr ¼ ai,Na 86 Sr 86 Sr river Na river i Na i i (15)

CO2 UPTAKE IN THE HIMALAYAS AND THE TIBETAN PLATEAU

where X stands for Ca, Mg, HCO3, Cl and Sr; i indicates the four source reservoirs (rain, evaporite, carbonate, silicate); and ai,Na is the mixing proportion of Na from the four reservoirs. Starting from an a priori set of end-member compositions, (X/Na)i, the equations are iteratively solved for ai,Na and (X/Na)i. In this case, an appropriate choice of a priori end members is critical.

Results and discussion Ternary diagrams and overall chemical composition Ternary diagrams are useful for showing the raw data representing different river stages, because these normalize for different discharge and allow comparison of relative elemental ratios. At the pH of river waters, alkalinity can be taken as bicarbonate (HCO2 3 ) concentration. For all river systems there is a large population of data centred around the Ca:Mg ratio of 2:1– 3:1 and stretching toward the Na þ K apex on the cation ternary, and around the alkalinity apex stretching to the Cl þ SO4 apex on the anion ternary. This suggests a predominance of carbonate and minor evaporite dissolution (Fig. 2). The Ganges tributaries, especially the Ghaghara and Kosi, and the Brahmaputra have noticeably higher Si proportions, indicating silicate weathering. Alkalinity (HCO2 3) can be generated by both carbonate and silicate weathering [reactions (1) and (2)]. Forward and inverse models are used to quantify the silicate weathering fraction from this me´lange of mainly carbonate and evaporite contributions (see section ‘silicate fractions’). There is quite a range in the Ca:Mg ratios observed. This is a measure of calcite:dolomite or felsic:mafic ratios of the substrate being drained.

Strontium isotope ratios With a box-and-whisker plot, the frequency distributions of 87Sr/86Sr in each basin can be clearly demonstrated. The line in the centre of the box is the median, and the box is defined by the lower and upper quartiles. The whiskers indicate minimum and maximum points excluding the outliers. The outliers are points that fall above the ‘upper quartile þ 1.5  interquartile difference’ or below the ‘lower quartile – 1.5  interquartile difference’. The large rivers of the eastern Tibetan Plateau, the streams of the Plateau interior as well as the rivers draining the two Himalayan syntaxes all have median 87Sr/86Sr values less than 0.72 (Fig. 3). The outlier streams of the Indus can be as radiogenic as 0.8448, but among the remainder 87Sr/86Sr above

135

0.72 are rare. Only the Ganges tributaries of the Himalayan front, especially the Ghaghara, have median values greater than 0.74. Maximum 87Sr/ 86 Sr ratios are 1.244 for the Alaknanda-Bhagirathi of the Ganges source, 1.022 for the Ghaghara, and 1.064 for the Kosi headwater tributaries. The 87Sr/86Sr v. 1/Sr plot is useful for demonstrating the mixing relationships between the Sr of carbonate origin with high Sr concentrations and typically unradiogenic 87Sr/86Sr c. 0.709 and the Sr of silicate origin with lower Sr concentrations and variable isotope ratios depending on the Rb/Sr content and the age of rocks being weathered. Data for different river systems are plotted on the same scales to accentuate the inter-basin differences (Fig. 4). A large number of data points in most drainage basins occupy the high Sr concentration corner with 87Sr/86Sr c. 0.71 (Fig. 4). The data become scattered at lower Sr concentrations because of the diversity of silicate rocks in terms of age and Rb/Sr composition. As was seen in the boxand-whisker plot (Fig. 3), the Ganges tributaries are radiogenic, and the Sr concentrations of many of the radiogenic samples are quite high. In the Indus, the bulk of the headwater samples have low Sr isotope ratios, and radiogenic values are limited to the right bank tributaries of the Indus in the very headwaters. The large flux of highly radiogenic Sr of the Ganges is not as prevalent in the Indus valley or in the upper Brahmaputra, the two Himalayan syntaxes. The Ganges with its high 87Sr/86Sr and high Sr concentrations played an important role in the evolution of the Sr isotopic composition of the ocean in the Cenozoic era. The source of the radiogenic Sr has been variously ascribed to granites/gneisses (Edmond 1992; Krishnaswami et al. 1992; Pande et al. 1994; Singh et al. 1998), metamorphosed carbonate rocks (Blum et al. 1998; Harris et al. 1998; Palmer & Edmond 1992; Quade et al. 1989), and metasediments (Harris 1995). The exact source of radiogenic Sr in the Ganges tributaries is difficult to quantify and is still a contentious issue, but Bickle et al. (2003), based on a large number of Sr and isotope data of the tributaries of the Ganges source waters (the Alaknanda), argued that all four major litho-tectonic units contribute significantly to the input of radiogenic Sr in the Ganges headwaters with perhaps the HHCS and the LHS playing a more significant role. Although it is not my intention to specify the exact source of the radiogenic Sr in the Ganges source waters, the analyses of data for the HTP suggest that this feature is localized to the Himalayan front draining the HHCS and LHS and is not observed in other peripheral regions of the Tibetan Plateau or in the two syntaxes. If we plot only the Ganges data on an 87Sr/86Sr v. Ca/Sr plot (Fig. 5), the spread in the data set can be

136

Table 2. Weathering fluxes for representative tributaries of the Himalayas and the Tibetan Plateau and rivers in other orogenic zones River

Salween Mekong Chang Jiang (Yangtze) Chang Jiang (Yangtze) Hong (Red) Huang He (Yellow) Indus

Location

fCO2 105 mol km22 a21

(Ca þ Mg)carb SCatsil (Ca þ Mg)sil

(Ca þ Mg)sil

Month

534 421 294

49.1 49.3 76.4

92.0 117 260

8.02 10.3 21.0

48.0 17.3 13.0

42.5 51.1 82.4

15.2 16.4 36.8

3.8 4.0 9.0

0.41 0.34 0.35

8 1, 9 8

Dadu He

698

47.2

67.6

6.58

4.8

38.4

9.5

2.2

0.33

5

212 172

4.4 39.9

20.8 232

3.30 15.3

19.6 16.0

17.5 70.4

5.3 29.3

0.6 5.7

0.30 0.24

9 5– 6

132

23.2

176

3.13

35.1

12.9

15.5

5.8

0.88

summer

1125

10.8

9.6

1.48

23.5

11.5

7.0

2.1

2.19

9

1064

8.3

7.8

0.86

39.0

3.35

3.3

1.0

1.22

9

1195

14.1

11.8

1.49

23.4

7.30

4.4

1.2

1.03

9

1114

13.7

12.3

2.54

9.3

14.9

2.3

0.5

0.38

1553

49.4

31.8

7.89

13.0

42.9

10.2

2.1

0.67

583 145

51.3 29.7

88.0 205

4.15 4.51

11.0 46.0

25.6 82.9

5.7 37.2

1.6 10.3

0.18 0.50

49.8

0.64

Narayani @ Narayanghat Ganga – Kosi Arun ab Sun Kosi Brahmaputra Lower Tsangpo @ Pasighat Sum/Average of HTP Other orogenic zones Yukon Pelly Stewart Yukon @ Carmacks Mackenzie Liard @ Liard River Athabasca @ Hinton

660

467

1331

90.5

20.8

504

198

Reference

Noh et al. 2009 Noh et al. 2009 Noh et al. 2009 Ellis et al. unpublished Moon et al. 2007 Wu et al. 2005 Karim & Veizer 2000 Dalai et al. 2002a Krishnaswami et al. 1999 Bickle et al. 2003

6, 8, 9, 11 Galy & France-Lanord 1999 6, 8, 9, 11 France-Lanord et al. 2003 11 Oliver et al. 2003 4 Singh et al. 2005

251 374 292

12.3 13.1 23.9

49.0 35.0 81.8

2.31 1.98 0.79

10.9 6.9 14.7

14.3 20.1 14.5

2.3 2.0 3.3

0.6 0.5 0.8

0.11 0.14 0.10

8 8 8

Millot et al. 2003 Millot et al. 2003 Millot et al. 2003

351

11.7

33.4

1.55

7.5

14.5

1.6

0.4

0.11

8

Millot et al. 2003

558

5.5

9.8

0.42

5.2

0.4

0.1

0.10

6

Millot et al. 2003

5.76

Y. HUH

Salween ab. Daojie Mekong @ Da Hai Jinsha @ Geli Ping

Red @ Yuan Jiang Huang He @ Lanzhou Indus @ Thakot Bridge Ganga – Yamuna Yamuna @ Batamandi Ganga Source Bhagirathi @ Devprayag Ganga Source Alaknanda @ Devprayag Ganga – Ghaghara Bheri @ Sampujighat Ganga – Gandak

fCO2 109 mol a21

Runoff* Discharge Area TDS SCatsil/ mm km3 a21 103 km2 106 t a21 SCattot mole %

Amazon

Yana Indigirka Global

1200

34.7 175

69.9

3.47

7.2

146

80.8

16.7

33.5 242

3.5

0.8

0.12

6

Millot et al. 2003

80.8

19.6

1.35

7

Stallard & Edmond 1983 Stallard & Edmond 1983 Stallard & Edmond 1983 Stallard & Edmond 1983 Edmond et al. 1996 Edmond et al. 1996 Edmond et al. 1996 Edmond et al. 1996 Edmond et al. 1996 Edmond et al. 1996 Edmond et al. 1996 Edmond et al. 1996 Huh et al. 1998 Huh et al. 1998 Huh et al. 1998 Huh et al. 1998 Huh et al. 1998 Huh et al. 1998 Huh et al. 1998 Huh et al. 1998 Huh et al. 1998 Huh et al. 1998 Huh et al. 1998

400

41.6

104

12.3

29.7

19.9

12.3

2.9

0.28

6

600

74.9

125

25.0

42.4

24.3

25.0

6.2

0.50

6

1000

83.3

83.3

16.4

9.3

60.1

1.6

3.3

0.40

12

1000

1.5

1.5

0.44

73.2

0.10

0.4

0.1

0.72

10

1000

1.2

1.2

0.25

69.7

0.08

0.3

0.1

0.50

10

Caparo @ Rt. 5

800

0.8

1.0

0.16

36.2

0.20

0.2

0.0

0.40

10

Portuguesa @ Rt. 5

800

0.7

0.8

0.26

37.0

0.32

0.3

0.1

0.82

10

Guanare @ Rt. 5

1000

1.8

1.8

0.68

20.6

1.90

0.7

0.2

0.85

10

Cojedes @ Rt. 5

300

0.6

2.1

0.56

14.3

2.15

0.6

0.1

0.65

6

Acarigua @ Rt. 5

800

2.0

2.5

0.48

10.8

2.91

0.5

0.1

0.51

10

Morador @ Rt. 5

800

0.7

0.8

0.24

24.4

0.54

0.2

0.1

0.76

10

Barayy Tukulan Kelly Allakh-Yun Khanda Tyry East Chandyga Tompo Tuostakh Moma Indigirka ab. Moma

150 150 199 200 200 200 200 200 100 200 199

0.5 0.4 2.1 5.0 1.6 2.7 2.0 8.4 2.1 7.0 28.1

3.6 2.8 10.7 24.8 8.2 13.7 9.9 41.8 20.6 35.0 141

0.06 0.04 0.23 0.44 0.10 0.27 0.06 0.71 0.20 1.11 4.55

16.6 11.7 15.8 11.3 3.1 7.3 2.1 10.8 11.8 23.9 38.7

0.21 0.19 0.87 2.59 2.28 2.46 2.23 4.42 1.16 2.54 5.21

0.1 0.0 0.2 0.4 0.1 0.3 0.1 0.7 0.2 1.1 4.6

0.0 0.0 0.1 0.1 0.0 0.1 0.0 0.2 0.0 0.3 1.1

0.04 0.03 0.05 0.04 0.03 0.05 0.01 0.04 0.02 0.08 0.08

8 8 8 8 8 8 7 8 8 7 7

Canagua @ Rt. 5

Aldan

497

37358

99259

550

8700

2500

CO2 UPTAKE IN THE HIMALAYAS AND THE TIBETAN PLATEAU

Orinoco

Peace @ Hudson’s Hope Ucayali @ Pucallpa Beni ab. Madre de Dios Madre de Dios ab. Beni Huallaga ab. Shanusi Paguey @ Rt. 5

Gaillardet et al. 1999

*The mean annual runoff for the Indus and Kosi was estimated by the author using GIS.

137

138

Y. HUH

Fig. 2. Cation (left) and anion (right) ternary diagrams for rivers of (a)(b) the eastern Tibetan Plateau and the Plateau interior, (c)(d) the Himalayan syntaxes, and (e)(f) the Himalayan front. The number of data plotted for each river system is in parentheses. The data sources are listed in Table 1.

explained by mixing of three end-members: silicate (Ca/Sr ¼ 250, 87Sr/86Sr c. 1.2), Precambrian carbonate rocks (Ca/Sr ¼ 5000, 87Sr/86Sr ¼ 0.715 + 0.01), and a third component probably evaporite, phosphate and carbonate rocks different from the Precambrian carbonate rocks considered above (Ca/Sr ¼ 200, 87Sr/86Sr ¼ 0.715). The carbonate end-member is based on the analyses of the Precambrian carbonate rocks of the Lesser Himalaya (Singh et al. 1998), and the evaporite end-member is from Krishnaswami et al. (1999). A continent—continent collision seems to be the ultimate mechanism whereby such radiogenic Sr

at high flux is observed, whatever its immediate source. If this is so, the radiogenic phases in the marine Phanerozoic 87Sr/86Sr record could indicate periods of Himalayan type orogeny.

Silicate fractions From the forward or inverse modelling, one can obtain contributions to the dissolved load from the pre-assigned reservoirs, for example, atmospheric, evaporite, carbonate, and silicate. In this paper, the main interest is with the draw-down of atmospheric CO2 and the emphasis will be on the fraction

CO2 UPTAKE IN THE HIMALAYAS AND THE TIBETAN PLATEAU

139

Fig. 2. (Continued).

of cations from silicate weathering, SCatsil/ SCattot. The results are shown as box-and-whisker plots (Fig. 6). The rivers draining the eastern Tibetan Plateau have similar SCatsil/SCattot. The values for internally drained rivers may be overestimated, because some of the Nasil may actually be associated with sulphate salts, but in most forward approaches all sulphate is binned with Ca. The Indus and many of the tributaries of the Ganges and the Brahmaputra have higher median values of SCatsil/SCattot as well as larger ranges. Even within the Himalayan front, there are local differences: the Yamuna and the Alaknanda-Bhaghirathi, and Kosi systems are especially high and varied. The Ghaghara

and Gandak have similar ranges of values as the Tibetan rivers. Another interesting feature is that the river systems with radiogenic 87Sr/86Sr do not necessarily have high silicate fractions. Only the Yamuna, Alaknanda-Bhagirathi, and the Kosi of the Ganges have both radiogenic 87Sr/86Sr and high silicate fractions. The Indus and Brahmaputra have high silicate fractions but ordinary 87Sr/86Sr (Figs 3, 6). If the source of Sr in these samples is exclusively silicate rocks, there should be a good correlation between the 87Sr/86Sr ratio and SCatsil/SCattot. As can be seen in Figure 7, there is a broad correlation between the two for some but not the majority of samples. Plotting 87Sr/86Sr ratio

140

Y. HUH

Fig. 2. (Continued).

against (Na þ K)sil/SCattotal does not improve the correlation (not shown). Weathering of silicate rocks generally releases radiogenic Sr but at low flux and does not lead to dramatic increases in 87 Sr/86Sr in most cases.

Chemical weathering rates Many published studies do not report runoff or discharge values of the individual samples for which chemistry data were obtained. Unless one carries a current meter on field expeditions, an exact estimate of the discharge at the time and place that the samples are collected is impossible. Even if these

portable current meters were available logistically, the larger uncertainty with spot sampling precludes this as a useful method unless one is prepared to carry out a full time-series sampling. In some studies, the mean annual discharge or runoff was estimated by a GIS approach using global digital databases based on interpolated runoff from those measured at hydrological stations (Wu et al. 2005; Moon et al. 2007; Noh et al. 2009). Hren et al. (2007) used satellite measurements of rainfall to calculate rainfall-weighted flow accumulations. Due to the lack of runoff data for individual samples, the fluxes and area-averaged yield of silicate weathering and uptake of atmospheric CO2

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Fig. 3. Box-and-whisker plot of the strontium isotope ratios. There is a break in the y-axis between 0.85 and 1.00. The number of samples is indicated on the graph for each river system. The data sources are listed in Table 1.

can only be calculated for the representative major tributaries for which discharge data are available in the literature (Table 2). They are the furthest downstream samples among those considered here but are still in the HTP or just at the exit of the HTP (Fig. 1). To make comparisons, exit samples were chosen also for basins that do have runoff data for individual samples. Total chemical weathering rates are calculated as: fTDS ¼ TDS  mean annual runoff  basin area

(16)

where total dissolved solids (TDS) is Na þ K þ Mg þ Ca þ Cl þ SO4 þ CO3 þ SiO2. The TDS flux is a measure of chemical weathering rate irrespective of the specific rock type being weathered. The area averaged TDS yield is similar for the eastern Tibetan Plateau rivers and the Yamuna and the source waters of the Ganges (Fig. 8). The Ghaghara

and Gandak have higher values, but carbonate weathering is the chief contributor (Table 2). The dissolved load is only a minor part of the weathering products carried by rivers, and the suspended material constitutes the major component. The suspended load is notoriously difficult to estimate by spot sampling, because flash floods can deliver the bulk of the suspended material (Qin et al. 2006). There is one study that measured suspended sediment load twice daily along with water discharge at ten river-gauging stations in the Marsyandi basin, tributary to the Gandak, for five years (Gabet et al. 2008). The total erosion rate including bedload and solute load was on the order of 0.1–2 mm a21. Only about one tenth of this would be due specifically to the solute load. These rates estimated from material carried by rivers are instantaneous rates at the time sampled. On the other hand, cosmogenic isotopes (10Be and 26Al) and apatite fission track ages represent long term bedrock incision rates. The latter estimates range

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Fig. 4. The strontium isotope ratios of the rivers of (a)(b) the eastern Tibetan Plateau and the Plateau interior, (c)(d) the Himalayan syntaxes, (e)(f) the Himalayan front. They are plotted on the same scale to accentuate the inter-basin differences. The data sources are listed in Table 1.

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Fig. 5. The 87Sr/86Sr plotted against Ca/Sr molar ratios for only the Ganges tributaries. Potential end-members for three-way mixing are illustrated. S, silicate, C, carbonate, E, evaporite. The data sources are listed in Table 1.

between c. 1.5 mm a21 in the Central Nepalese Himalaya to 9–12 mm a21 in the western Himalayan syntaxis (Leland et al. 1998; Zeitler 1985). Of the TDS carried by 62 largest rivers of the world (2130  106 t a21) (Gaillardet et al. 1999), the HTP rivers studied here deliver 90  106 t a21, c. 4%, whereas the HTP occupies only about 1 % (1.33  106 km2) of the area (99.3  106 km2). It is commonly assumed that the strongest physical and chemical weathering rates are found in the HTP region, since it is the most conspicuous collisional regime in the Cenozoic era. This is mainly a consequence of Milliman & Meade’s (1983) compilation of sediment export rates, which estimated that at mouths, the HTP rivers (excluding the Salween and internal drainages) contribute 27% of global annual suspended sediments while occupying only 7% of the area. However, the dissolved flux is a minimal fraction of the total weathering rate. There is still a potential that the sediments generated in the Himalayas, transported to the alluvial plains and stored there, will undergo further weathering in a warmer and reaction-dominated system to generate high weathering rates. However, the average flux of TDS transport in the HTP portion of the Ganges is c. 41% (18  106 t a21) (this study, Table 2) of the entire Ganges basin (44  106 t21, Gaillardet et al. 1999). Since the headwaters

account for only c. 15% in terms of area, this is proportionally high input from the mountains and does not suggest significant input of TDS from the plains. Comparisons can be made to the rivers draining other orogenic zones of the world. Data from the Amazon and Orinoco draining the Andes, Mackenzie and Yukon draining the Rockies, and the Siberian rivers draining the mountains of the Russian far east are available in the literature (Edmond et al. 1996; Huh et al. 1998; Stallard & Edmond 1983). To make relevant comparisons, I hand-picked from the literature data set only those rivers draining the highlands and did not include samples that flow in the plains. From my best estimate, the average rate of TDS transport in the Andes is similar to that of the HTP (Fig. 8). However, because of the extremely low runoff, the rates are lower in the Rockies and Siberian mountains.

Consumption rate of atmospheric CO2 Three types of CO2 consumption rates are reported in Table 2. First, the CO2 taken up during weathering of CaCO3 (fCO2 carb), which will have no net effect on million-year timescales, is 504  109 mol a21 for the HTP system studied here. This is c. 4% of the global flux (12 300  109 mol a21). Second is the CO2 taken

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Fig. 6. Box-and-whisker plot for SCatsil/SCattot values. The number of samples used is indicated for each river system. The trend for S(Ca þ Ma)sil/SCattot is similar (not shown).

up during weathering of silicate minerals including the alkali-silicate minerals (fCO2 sil short-term). In total the HTP rivers (198 109 mol a21) consume c. 2% of the global total (8700  109 mol a21). Finally, the CO2 fixed during weathering of Caand Mg- silicate minerals and subsequent deposition of carbonate minerals (fCO2 sil long-term) is also about 2% of the global (2500  109 mol a21) for the HTP (50  109 mol a21). The HTP rivers studied here cover about 1% of total exhoreic drainage area and is about the same in terms of water discharge. The HTP is contributing disproportionately large quantities of dissolved material to the oceans but this is mainly due to carbonate and not silicate weathering. The area-averaged CO2 consumption rates (fCO2 sil long-term) are similar for the eastern Tibetan Plateau rivers, the Ghaghara tributary system of the Ganges, and the Brahmaputra at c. 35 103 mol km22 a21 (Fig. 9). Per unit area,

the Yamuna and Alaknanda-Bhagirathi of the Ganges and the Indus have more than twice as high CO2 consumption rates. The Kosi system displayed signs of extremely silicate dominant regime, for example high 87Sr/86Sr and SCatsil/ SCattotal. However, because of the low runoff, the CO2 consumption rates are not spectacular. The Indus also has high SCatsil/SCattotal but the precipitation and runoff in this region is also low. Thus, even within the HTP system, the tributaries of the Himalayan front in general have higher CO2 uptake rates. The HTP and other orogenic regions together are responsible for the drawdown of atmospheric CO2 by silicate weathering at present. In this respect, the HTP is not anomalous.

Uncertainties There is an inevitable logistical bias in sampling. For some rivers the extreme headwaters have

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Fig. 7. The relationship between 87Sr/86Sr and SCatsil/SCattot in the rivers of (a)(b) the eastern Tibetan Plateau and the Plateau interior, (c)(d) the Himalayan syntaxes, and (e)(f) the Himalayan front. The axes scales are all different to illustrate the details of the data set. The trend is similar when plotted against S(Ca þ Ma)sil/SCattot or S(Na þ K)sil/SCattot (not shown). The data sources are listed in Table 1.

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Fig. 8. Total chemical weathering rates estimated from the total dissolved solids (TDS) of (a) the rivers of the Himalayas and the Tibetan Plateau. Comparison can be made to other orogenic zones of the world: (b) the Amazon and (c) Orinoco draining the Andes, (d) Mackenzie and (e) Yukon draining the Rockies, and (f) Siberian rivers. Different bars within each river system indicate samples from different tributaries.

been studied in detail (e.g. the Kosi), and in others only the major tributaries have been sampled. This will generate some bias in terms of 87Sr/86Sr or SCatsil/SCattotal if the headwaters have radiogenic silicate rocks weathering at high intensity. In addition to the logistical sampling bias spatially, there is a temporal bias. Conventionally,

the samples are collected in the rising or falling stages, and the concentrations are multiplied by mean annual discharge to obtain the annual average flux. This is adequate to within c. 15% in basins where the lithology is largely similar (Qin et al. 2006). However, in basins where the changes in precipitation affect not only the total runoff but

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Fig. 9. The consumption rate of atmospheric CO2 for the samples from each system furthest downstream within the Himalayas and the Tibetan Plateau. For comparison other orogenic zones of the world are also shown as in Figure 8.

also the specific rock type that is being weathered, this can be a more serious problem. The few samples collected during the monsoon indicate that the SCationsil are lower due to dilution (Krishnaswami et al. 1999). Here lies the fundamental uncertainty with spot sampling. For example, Himalayan rivers run across different zones, the

TSS, HHCS and LH. Depending on the season, different parts may contribute more dissolved load to the Ganges main tributaries like the Yamuna, Ghaghara, and so on. West et al. (2002), based on small basins with well-characterized variations in hydrography and high temporal resolution sampling of the riverine chemistry, indicated that spot

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sampling can introduce large error. Bickle et al. (2003), based on modelling of Sr and 87Sr/86Sr of the tributaries and main channels, also showed that there are large variations in the relative fluxes from the different units (TSS, HHCS, LH) and that ‘single sampling campaigns do not provide good characterization of river chemical fluxes and their sources in the Himalaya’. Even aside from spot sampling problems, there is a major uncertainty in the runoff estimates. Although more recent publications use digital databases to obtain averaged values of each sample drainage basin, in earlier studies runoff values are available only for the larger rivers. Wolff-Boenisch et al. (2009) is perhaps the only study combining discharge and chemical measurements for the same locations. They estimate a 20% uncertainty in discharge even with station measurements. For those studies investigating specifically the sources of radiogenic Sr, estimation of relative water and chemical input fluxes using the change in the mainstream composition sufficed and circumvented the need for runoff data (Bickle et al. 2003; English et al. 2000). However, correct runoff estimates are essential for deriving rates of CO2 uptake. The (Ca þ Mg)sil/Nasil ratios used to calculate Casil and Mgsil are associated with a large uncertainty. They are usually based on local bedrock composition, soil profiles or small streams flowing only through silicate lithology. Using a single ratio to treat all samples within a river system fails to account for individual rock types in sub-drainage basins. For example, for the Ganges source waters draining the High Himalayas and the Lesser Himalayas, both are assigned the same ratios because of the difficulties of reliably assigning different values to them. The uncertainties of c. 35% in the ratios, when propagated, lead to c. 14% uncertainties in the SCationsil (Krishnaswami et al. 1999). Wolff-Boenisch et al. (2009) estimated a 60% error in the (Ca/Na)sil and (Mg/K)sil ratios based on the arithmetic mean of the extreme values. The sulphuric acid from oxidative weathering of sulphides, if used to dissolve silicate rocks, can introduce a positive bias on the CO2 consumption rates by silicate weathering. This is difficult to estimate in many cases, though sulphur isotope ratios may help eventually. Conventionally, upper and lower estimates are derived by assuming either that all sulphate is from evaporites or that all are from sulphide oxidation and that carbonate and silicate rocks are weathered in proportion to that being weathered by CO2. Rarely do the studies employ stable sulphur or carbon isotope ratios to derive a more quantitative estimate of the sulphide v. gypsum contribution. In this paper, only weathering in the Himalayas and the Tibetan Plateau has been considered.

Suspended sediment load may be stored in large alluvial plains and removed during periods of increased discharge during the monsoonal months. Thus, secondary effects of delayed weathering of HTP material has to be kept in mind.

Future work The CO2 consumption rates vary among rivers draining the Himalayas and the Tibetan Plateau. The eastern Tibetan Plateau rivers are not very radiogenic, their silicate contributions are low, and the CO2 consumption rates for short and long term are on the order of (0.6–7.3)  105 mol km22 a21 and (0.2– 2.2)  105 mol km22 a21, respectively. The two syntaxes of the Himalayas, the Indus and the Brahmaputra, have variable 87Sr/86Sr, the silicate contributions are higher than the eastern Tibetan Plateau rivers, but the total CO2 uptake rates are low because of lower runoff in the headwaters. The Ganges tributaries draining the Himalayan front can be extremely radiogenic, have higher median silicate contributions than the eastern Tibetan Plateau rivers, and the CO2 consumption rates are about twice as high. Exposure of silicate lithologies with radiogenic Sr and high runoff seems to be responsible. Even the high CO2 consumption rates of the Himalayan front are not anomalous globally, considering the uncertainties with spot sampling and calculation of silicate fractions. The rate of CO2 draw-down by silicate weathering in the Himalayan front rivers are on the same order of magnitude as other orogenic rivers of the world, although fluxes of 87Sr are higher. To constrain these rates better, one of the critical parameters is runoff. At regional scales, digital methods with regional coverage such as using satellite precipitation data to convert to runoff or adding more hydrological stations to the network of river discharge would provide better interpolations. Also, instead of discharge only at select locations, it would be useful to calculate runoff for each sample in order to observe factors regulating chemical weathering of silicate rocks in the basins. A digital conversion of environmental factors would be useful, foremost the lithology. The SCationsil is expected to vary consistently with the exposed lithology of the drainage basin (% silicate) (Krishnaswami et al. 1999), but this cannot be quantified because of lack of detailed digital lithological (as opposed to geological) maps. Du¨rr et al.’s (2005) map, if increased in resolution and made available to the community, would be helpful. Even were this available, there would still remain the uncertainty of formations that are hidden by lateritic cover but still reached by surficial water. Another is vegetation which has usually been assumed to co-vary with climate and substituted

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by runoff and temperature parameters. This has been due mainly to a lack of adequate method of quantifying the vegetation data. One method would be to use NDVI (Normalized Differential Vegetation Index) obtained from satellite data. The uncertainties driving both the forward and inverse methods of calculating the silicate percentage is the knowledge of local silicate rock compositions, for example (Ca/Na) or the (Mg/ Na) ratios, and how incongruently they weather. Some basic data acquisition of local bedrock or small silicate streams in these remote regions is required. The effect of sulphuric acid in dissolving silicate rocks is difficult to quantify. Developing a method to quantify this factor, perhaps using the double isotope technique of sulphate would be useful. The author thanks C. Burchfiel, K. Hodges, P. Molnar, W. Moore, M. Raymo, A. Stewart, N. Walton, and K. Whipple for sample collection and A. Ellis, S. Moon, H. Noh, and L. Wu, for helpful discussions. This work was supported by US NSF EAR 134966 and by the Korea Research Foundation Grant funded by the Korean Government (MOEHRD, Basic Research Promotion Fund) (KRF-2007-531-C00054).

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consumption and Sr flux. Chemical Geology, 234, 308– 320. Stallard, R. F. 1980. Major element geochemistry of the Amazon River system. Ph.D. thesis, MIT-WHOI Joint Program in Oceanography. Stallard, R. F. & Edmond, J. M. 1983. Geochemistry of the Amazon 2. The influence of geology and weathering environment on the dissolved load. Journal of Geophysical Research, 88, 9671– 9688. Summerfield, M. A. & Hulton, N. H. 1994. Natural controls of fluvial denudation rates in major world drainage basins. Journal of Geophysical Research, 99, 13871– 13883. White, A. F. & Blum, A. E. 1995. Effect of climate on chemical weathering in watersheds. Geochimica et Cosmochimica Acta, 59, 1729–1747. Wolff-Boenisch, D., Gabet, E. J., Burbank, D. W., Langner, H. & Putkonen, J. 2009. Spatial variations in chemical weathering and CO2 consumption in Nepalese High Himalayan catchments during the monsoon season. Geochimica et Cosmochimica Acta, 73, 3148– 3170. Wu, L., Huh, Y., Qin, J., Du, G. & Van der Lee, S. 2005. Chemical weathering in the Upper Huang He (Yellow River) draining the eastern Qinghai-Tibet Plateau. Geochimica et Cosmochimica Acta, 69, 5279– 5294. Zeitler, P. K. 1985. Cooling history of the NW Himalaya, Pakistan. Tectonics, 4, 127–151. Zeitler, P. K., Chamberlain, C. P. & Smith, H. A. 1993. Synchronous anatexis, metamorphism, and rapid denudation at Nanga Parbat (Pakistan Himalaya). Geology, 21, 347–350.

Evolution of the Indian summer monsoon: synthesis of continental records PRASANTA SANYAL1 & R. SINHA2* 1

Department of Geology and Geophysics, Indian Institute of Technology, Kharagpur, India 2

Department of Civil Engineering, Indian Institute of Technology Kanpur, Kanpur 208016, India *Corresponding author (e-mail: [email protected])

Abstract: Fluvial sediments of the Siwalik succession in the Himalayan Foreland Basin form the most important continental archive for reconstructing monsoonal fluctuations during the Late Miocene to Late Pleistocene. A number of proxy records suggest multiple phases of monsoonal intensification with peaks at 10.5, 5.5 and 3 Ma after which the strength of the monsoon decreased to modern day values with minor fluctuations. Detailed evaluation of Late Quaternary interfluve stratigraphic development in the Ganga plains shows that interfluve areas near the major rivers aggraded periodically between 27 and 90 ka. They subsequently degraded or accumulated sediment only locally, probably reflecting decreased monsoonal precipitation around the Last Glacial Maximum. Increased precipitation during the 15 to 5 ka period of monsoon recovery probably increased discharge and promoted incision and widespread badland formation. In western India, the fluvial records back to c. 128 ka suggest a stronger monsoon around 80 ka followed by periods of weakened monsoon around 70 to 30 ka and progressive desiccation in the glacial period. Holocene lacustrine records from western Rajasthan suggest maximum lake levels at c. 6 ka and complete desiccation between c. 3 and 4 ka.

The Indian monsoon is an important element of the climate system in the tropics. A seasonally reversing wind, which characterizes the Indian monsoon, brings moisture from the Arabian Sea and the Bay of Bengal during summer (June to September) (Fig. 1) and causes rain in most parts of the Indian subcontinent. In contrast, air blows from land during the winter season resulting in a dry winter monsoon. It has been suggested that the monsoon system was initiated about 20 Ma ago because of uplift of the Tibetan Plateau beyond a critical height (Harrison et al. 1992; Prell & Kutzbach 1992; Molnar et al. 1993). During summer, heating of the Tibetan plateau creates low atmospheric pressure, which acts as a powerful pump for moist air from the oceans, resulting in heavy rain (Quade et al. 1995). Reverse circulation of wind occurs during the winter; the radiative cooling of Tibetan Plateau causes flow of cold dry continental air towards the Indian Ocean. Experiments with atmospheric general circulation models have shown that changes in elevation of Himalaya-Tibet have significant effects on the intensity of monsoon (Prell et al. 1992; Clemens & Prell 2003). Simulations with no mountains or reduced elevation result in significantly weaker or even no monsoonal circulation. Monsoon intensity could also be affected by changes in surface boundary conditions, such as the albedo of Africa –Asia

landmass, the extent of snow cover over Tibet, the sea surface temperature of the Indian Ocean and the concentration of atmospheric CO2 (Prell et al. 1992). A reasonably good record of variation of monsoon intensity through time has been obtained from oceanic sediments (Kroon et al. 1991; Nigrini 1991; Prell et al. 1992; Prell & Kutzbach 1992). During the summer monsoon, the wind flow from the Indian Ocean causes transport of surface water and develops intense centres of upwelling in the ocean. The upwelling brings cold, nutrient-rich waters from several hundred metres depth to the surface and triggers high productivity in the photic zone (Swallow 1984; Anderson & Prell 1993). The duration and intensity of the upwelling reflects the intensity of the monsoonal winds. Thus, the biological, chemical and sedimentological record of the upwelling system has a direct link to the structure and intensity of the monsoonal winds. Monsoonal variability from oceanic records is well constrained in terms of wind strength. In the Arabian Sea, intense seasonal upwelling is induced by the southwesterly monsoon winds (Anderson & Prell 1993). Sediments in the NW Arabian Sea exhibit a characteristic fauna (radiolarians and foraminifers) that are endemic to areas of upwelling. These biota are normally encountered only in cool temperate water, and therefore, their appearance and abundance in

From: Clift, P. D., Tada, R. & Zheng, H. (eds) Monsoon Evolution and Tectonics –Climate Linkage in Asia. Geological Society, London, Special Publications, 342, 153–183. DOI: 10.1144/SP342.11 0305-8719/10/$15.00 # The Geological Society of London 2010.

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Fig. 1. Wind pattern and depression track during Indian summer monsoon (after Sengupta & Sarkar 2006). Grey arrows indicate the depression track from Bay of Bengal and Arabian Sea. The thin dashed lines with small arrow indicate the direction of surface wind. Thick dashed lines indicate the date of onset of summer monsoon and the thick solid lines indicate the withdrawal of summer monsoon in different parts of India. The mixing of the Arabian Sea and Bay of Bengal vapour occurs near Allahabad.

sediments should indicate upwelling conditions. Miocene to Recent sediments from the NW Arabian Sea show distinct geochemical and biological changes, which suggest that monsoonal upwelling conditions were established about 8 Ma (Fig. 6c; Kroon et al. 1991). Pelagic sediments deposited before 10.5 Ma contains nanofossils, which are characteristic of warm water and relatively low productivity. Opal rich sediments, which reflect initiation of strong monsoon circulation, were deposited between 10.5 Ma and 8.0 Ma (Prell et al. 1992). Fluvial sediments of the Siwalik succession in the Himalayan foreland basin, which stretches from Burma in the east to Afghanistan in the west form the most important continental archive for reconstructing monsoonal fluctuations during Late Miocene– Pleistocene. This foreland basin is synsedimentary in nature, where sediments

deposited during the last 20 Ma form a pile as thick as 5 km in some places (Johnson et al. 1985). Post-Siwalik fluvial deposition in the Himalayan foreland basin has continued through the Quaternary without any break (Wadia 1932; Raiverman 2003; Sinha et al. 2007a); however, climatic records during this transition period are generally poor except for some studies in the intermontane valleys (Suresh et al. 2007; Sinha et al. 2009) which have described the Quaternary successions lying directly on the Siwaliks. It is believed that sedimentation in Siwalik basin stopped around 200 ka (Ranga Rao 1973; Singh et al. 2001) as a result of tectonic activity along the Himalayan Frontal Thrust (HFT) and associated folding, which led to the formation of an intermontane basin in the northwestern Himalaya and sedimentation continued as fan-terrace systems until the Late Holocene.

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A number of proxy records including palaeosols (Rettallack 1995; Thomas et al. 2002), soil carbonates (Quade et al. 1989; Sanyal et al. 2004, 2005a, b) pollens (Hoorn et al. 2000), microfossils (Phadtare et al. 1994), palaeomagnetism (Sangode & Bloemendal 2004) and general sedimentation pattern have been used to decipher the changes in the Indian summer monsoon from the Siwalik sediments. For the Quaternary period, continental records from northern and western India have been studied extensively and intensively as response systems to shifts in the monsoonal regimes, both in terms of magnitude and in the zone of influence. Major areas of interest include the Himalaya, the Ganga plains, the Thar desert, and the southwestern (Gujarat) and northeastern margins (Haryana) of the Thar desert (see Fig. 2 for locations of areas discussed in this paper). A variety of archives such as peat, loess, alluvial sediments, playas, lake sediments, dunes, tree-rings, and soil carbonates have been used to generate climate proxy data. This synthesis is focused on the continental records for understanding the evolution of the Indian monsoon covering two major time periods, Late Miocene to Late Pleistocene (written by P. Sanyal) and Late Quaternary (written by R. Sinha), and emphasizes the need for generating high resolution records

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from continental settings. The review is not intended to be comprehensive and the individual authors take full responsibility for the data presented in the respective sections.

Geological settings and stratigraphy The Siwalik sediments, which are part of the Himalayan foreland basin were deposited in the foredeep created during the Himalayan orogeny. The Siwalik sediments represent overfilled or last stage of foreland basin development where rate of sedimentation is greater than the rate of subsidence (Flemings & Jordan 1989). The sediments are dominantly fluvial and were probably deposited by an ancestral Ganga River system. These sediments are exposed in the southern flank of the Himalaya in a WNW to ESE trending belt (Fig. 2). They are bounded by the Main Boundary Thrust in the north and the Himalayan Frontal Thrust in the south. Near the main boundary thrust, the Siwalik Group are folded and faulted, but grade southward into flat beds overlain conformably by modern alluvium of the Himalayan foreland basin (Karunakaran & Rao 1979). They are characterized by alternating sandstone and mudstone, and conglomerate and

Fig. 2. Geographical extension of the Siwalik sediments in Himalayan foreland (after Gansser 1964) and locations of Siwalik sections described in this paper. Locations of cliff sections/cores in the Ganga plains, western Indian rivers and playas of Thar Desert described in this paper are also shown in boxes. Major thrust zones of Himalayas: ITSZ, Indus-Tsangpo Suture Zone; HFT, Himalayan Frontal Thrust; STDS, South Tibetan Detachment System; GCT, Great Counter Thrust; MCT, Main Central Thrust; MBT, Main Boundary Thrust.

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mudstone beds. The beds generally dip north or northeastward. The Siwalik basin is divided into a number of sub-basins separated from each other by lineaments (Virdi 1979; Raiverman et al. 1983). These lineaments are extensions of basement features extending from the Indian Shield to the Himalayas. They formed as normal faults during an extensional regime and were later reactivated as thrust faults during the Cenozoic orogeny (Dubey 1997). These faults not only controlled the thickness of the sedimentary successions, but also affected the sedimentation pattern (Raiverman et al. 1983). The thrust belts are characterized by sinuous traces with alternating re-entrants and salients. In this paper, we present data from the Ranital and Kotla section of the Kangra re-entrant, the Haripur Khol section of the Subathu salient, and

the Mohand Rao section of the Dehra Dun re-entrant, the Surai Khola section of Nepal and several sections (Chinji Nagri, Kaulial. Jalalpur, Rohtas, Mirpur and Pabbi Hills sections) from the Potwar Plateau of Pakistan (Fig. 3). The ages of the sedimentary successions in different Siwalik sections have mostly been determined by means of magnetostratigraphy (Fig. 4) (Johnson et al. 1982a, b, 1985; Appel et al. 1991; Sangode et al. 1996, 1999, 2003; Kumar et al. 1999, 2003a, b; Kumaravel et al. 2005). The magnetic reversal stratigraphy was determined using mudstones and fine grained sandstone from different Siwalik sections. This framework was correlated with the Geomagnetic Polarity Time Scale of Cande & Kent (1995) to obtain ages of the samples. Along with the stratigraphy based on palaeomagnetic normal-reversal sequence, magnetic fabric

Fig. 3. Simplified geological map of Indian sections in the Himalayan Foreland Basin (after Raiverman et al. 1983). Himalayan Foreland Basin is characterized by alternate salients and re-entrants. For the present study samples from Ranital and Kotla sections from Kangra re-entrant, Haripur Khol section from Subathu salient, Mohand Rao section from Dehra Dun re-entrant are chosen. (SRT, Salt Range Thrust; HFT, Himalayan Frontal Thrust; GTF, Ganga Tear Fault; YTF, Yamuna Tear Fault; MBT, Main Boundary Thrust).

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Fig. 4. Correlation of magnetic stratigraphy of (a) Kotla, (b) Ranital (both in Kangra), (c) Mohand Rao (Dehra Dun), (d) Haripur Khol (Subathu) and (e) Surai Khola section with modified GPTS of Cande & Kent (1995). The age ranges (in Ma) for the sections are: Kotla & Ranital c. 11– c. 6, Mohand Rao c. 10–c. 5, Haripur Khol c. 6– c. 0.5, Surai Khola c. 13– c. 1. The base of Ranital and Kotla section is demarcated by Jwalamukhi Thrust, Mohand anticline for Mohand section, 289 Dhanaura anticline for Haripur Khol section. The base of Surai Khola section was determined based on lithological 290 characteristics. The scale refers to the thickness of all the stratigraphic successions. (Source: Sangode et al. (2003); Sanyal et al. (2004, 2005a, b)).

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correlation, biostratigraphy, iterative matching, and conformable sedimentation rates in adjacent sections were also considered to constrain depositional ages. In addition, absolute ages obtained from dating of volcanic ash were used as a benchmark for correlation with geomagnetic polarity timescale. The Siwalik Group has been divided into three lithostratigraphic units (Pilgrim 1910, 1913): Lower (age: c. 18.3 to 10 Ma), Middle (c. 11 to 5.3 Ma) and Upper Siwaliks (5.3 to 0.22 Ma). The Siwalik Group, well known for its abundant vertebrate fossils, shows a distinct coarsening upward from mudstone-sandstone (Lower Siwalik), to sandstone-dominated (Middle Siwalik), to conglomerate, sandstone and mudstone (Upper Siwalik) facies. The Upper Siwalik succession shows significant lateral variability across the foreland basin, exhibiting an increase in conglomerate towards the Main Boundary Thrust (MBT). The Lower Siwalik and older strata are typically well indurated, whereas the Middle and Upper Siwalik strata are normally friable. The Siwalik Group is overlain by Quaternary conglomerates in broad synclines, as in the Dehra Dun Re-entrant of the foothills, and is also presumed to underlie the younger Quaternary strata of the Indo-Gangetic Plain to the south. Ages of sedimentary successions in the Kangra Sub-basin (Ranital and Kotla section, Fig. 4a, b) range from c. 11 to c. 6 Ma, in Haripur Khol from c. 6 to 0.5 Ma (Fig. 4d) and in Mohand Rao section c. 10 to 4.4 Ma (Fig. 4c). A 5 cm thick bentonized tuffaceous bed within overbank facies was discovered in Haripur Khol section by Kumar et al. (1999) at 1650 m level from base of the section. Fission track dating of zircons separated from lateral extension of this ash bed in an adjacent section (Tandon & Kumar 1984) gave an age of 2.14 + 0.5 Ma (Ghaggar River section in Punjab Sub-Himalaya, Mehta et al. 1993), which constrained its magnetic polarity to be benchmarked near the Gauss/Matuyama event. Details of magnetic polarity events, methods of deriving the ages and sedimentation patterns of various levels are described in several papers and reports of Sangode et al. (1996, 1999, 2003) and Kumaravel et al. (2005). The age of the Surai Khola section was first determined by Appel et al. (1991) using magnetic normal and reversal data. The presence of fossils was taken as tie points. The ages reported by Appel et al. (1991) were corrected (Sanyal et al. 2005b) based on the Geomagnetic Polarity Scale of Cande & Kent (1995) using the publication of Gautam & Ro¨sler (1999) (Fig. 4e). Age of sections in the Potwar Plateau has also been determined by palaeomagnetic methods. In addition, the absolute dating of volcanic ash beds present in these sections provided bench-marks

for calculating the sedimentation ages (Johnson et al. 1982a, b, 1985). However, high-resolution magnetostratigraphic age data from the same sections by Behrensmeyer et al. (2007) has been considered here. Because most of the data from Siwalik presented in this paper are from the different proxies collected from palaeosols, a brief description of palaeosol facies is given below.

Palaeosol facies Palaeosols in the Siwalik Group were mostly developed in floodplain deposits and vary in thickness from a few tens of centimetres to several metres. In exposed sections, palaeosols appear red, reddish-brown, grey and green, and are invariably mottled. Soil structure is extensive in the illuvial layers (B-horizons) of palaeosols (Sanyal et al. 2004). Most of the palaeosols contain carbonate in disseminated and nodular form. In the Kangra Sub-basin, palaeosols are grey, brown and yellow in colour and contain iron and calcareous nodules (Fig. 5), as well as root traces and biotubes. Green mottling is common in the palaeosols. Brown, purple and red palaeosols are common in the time range 11 to 7 Ma, but from 7 to 6 Ma yellow palaeosols dominate. Palaeosols are not observed in rocks younger than c. 6 Ma. In the Mohand Rao section around Dehra Dun (Fig. 4c), palaeosols are scanty but where available, they occur as discontinuous lenticular bodies within multi-storied sandstone. The palaeosols are mostly grey in colour but are locally brown. Green mottling and nodular carbonates are common in palaeosols. Locally mudstones show immature soil profiles (Kumar et al. 2004). Palaeosols show evidence of biological activities in the form of vertical, unlined burrows and surface traces. Thomas et al. (2002) studied palaeosols in the Haripur Khol section (Fig. 4d) in the northwestern Himalaya in detail. About 70% of the palaeosols are calcareous in nature and contain significant amounts of calcite in the form of nodules or fine calcareous matter disseminated in the groundmass. The colour and maturity of the palaeosols change from the bottom to top of the section. Red, moderate to strongly developed palaeosols grade into less developed yellow palaeosols starting at about 2.6 Ma. The lower part of the section is characterized by strong illuviation of clay (Bt horizons) in the palaeosols. Palaeosols are abundant in the Siwaliks of the Potwar Plateau area in Pakistan. Most of the palaeosols are characterized by a clayey subsurface (Bt Horizon). Similar to the Indian Siwalik sections, the palaeosols in the Potwar Plateau are extensively bioturbated and contain soil carbonate in the Bk Horizon. Some iron nodules are also present.

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Fig. 5. Field photograph showing a palaeosol with soil carbonate nodule (marked by the circle) in the Kangra sub-basin, NW Himalaya.

Various Siwalik sections are exposed in Nepal along road cuts, as well as in river beds. The best studied section is that at Surai Khola (Fig. 4e) in western Nepal. The colours of the palaeosols in Surai Khola section are red, yellow and greyishgreen. These palaeosols are characterized by the presence of an organic rich A-Horizon in most cases and nodular soil carbonate in the Bk Horizon. Red palaeosols are abundant in the lower part of the section and in the middle part the proportion of greyish-green palaeosol increases. In the upper part, yellow palaeosols with Mn-oxide nodules are abundant (Quade et al. 1995).

Monsoonal reconstruction from Siwalik successions using multi-proxy data Oxygen isotope ratio of soil carbonates In Siwalik palaeosols, soil carbonates occur in disseminated and in a nodular forms. To reconstruct monsoonal rainfall, nodular soil carbonates are considered as a closed system with little or no exchange with external fluids, and with little diagenetic alteration. The nodular soil carbonates are not ubiquitous within Siwalik successions. For example, in the Kotla and Ranital sections (Fig. 4a, b), soil carbonates are observed from 11 to 5.5 Ma and in the Haripur Khol section from 6 to 1.8 Ma. At Mohand Rao, soil carbonates are recorded only in the lower (9 to 8 Ma) and middle parts (5.5 to 4.5 Ma) of the section. As these four sections are close to each other, the oxygen isotopic data from

all sections along with the age of the samples have been combined to make a composite plot representing regional time variation of oxygen isotopes (Fig. 6c). In the composite plot, the oxygen isotope ratio shows three evolutionary phases. At around 10.5 Ma, d18O value is characterized by a much lower value of about –10‰; d18O tends to move towards a more positive value with decreasing age, reaching –6.6‰ at around 6.5 Ma. Subsequently, the isotope ratio is characterized by a sharp lowering, dropping to –9‰ at 6.0 Ma and then a second phase of higher values reaching – 6.5‰ at around 2 Ma. These indicate two phases of lower d18O values occurring around 10.5 Ma and 6.0 Ma punctuated by a period of maximum enrichment at around 6.5 Ma (Sanyal et al. 2004, 2005b). In the Potwar Plateau, soil carbonate nodules are available in sediments younger than 18 Ma (Quade et al. 1989, 1995) (Fig. 6a). Although d18O values are characterized by scattering, a closer look at the data reveals three distinct phases of d18O values. At 18 to 9.5 Ma, the d18O value averaged around – 10 to –9‰ with the lowest value of –11.5‰ observed at 9.5 Ma. Subsequently, the d18O value increases and the most enriched value was observed at 6 Ma. Post-6 Ma data are characterized by high scattering with d18O value ranges from –9 to –4.5‰. Oxygen isotope ratios in soil carbonate from the Surai Khola section of Nepal were studied by Quade et al. (1995) (Fig. 6b). Although the data are highly scattered, the depleted 18O/16O ratio around 10 Ma and enrichment of 18O/16O ratio subsequent to 6 Ma

160 P. SANYAL & R. SINHA Fig. 6. Oxygen isotope ratio variation in soil carbonate and pedogenic clay minerals in different Siwalik sections and comparison of monsoon with ocean record. (a) – (c) represent the d18O value of soil carbonate from Potwar Plateau of Pakistan, Surai Khola of Nepal, and Kangra (Ranital and Kotla)-Haripur Khol and Mohand Rao from Indian Siwalik respectively; (d) and (e) represent d18O value of pedogenic clay minerals from Potwar Plateau and abundance variations of G. Bulloides from Arabian sea respectively. The shaded parts represent time of monsoon intensification. Note that the monsoon intensified around 10.5, 5.5 and 3 Ma. (Sources: Quade et al. 1989, 1995; Kroon et al. 1991; Sanyal et al. 2004, 2005b).

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is consistent with the oxygen isotope ratio variations with time from other part of the Siwalik Group.

Oxygen isotope ratio of carbonate cement of sandstone In many Siwalik sections, soil carbonate nodules are not available which constraints our ability to reconstruct climatic events from different locations. Oxygen isotopic ratios of carbonate cement in sandstone can be considered as an alternate proxy to deciphering climate provided the late diagenetic imprints on the pristine isotopic ratio can be separated. An attempt was made to study the isotopic ratio of carbonate cement from the Surai Khola section of the Nepalese Siwalik, and compared with the isotopic ratio of soil carbonate from the same section to make a reference data set for comparison. The d18O values of carbonate cement of sandstone from the Surai Khola section also show three major evolutionary trends as a function of time (Sanyal et al. 2005a). From 12 to c. 6 Ma, the average d18O (VPDB) value is –13.6 + 1.9‰ (n ¼ 114) with a large spread from –10 to –18‰ (Fig. 7). From 6 to 4 Ma, d18O values show a sudden swing towards higher values with less scatter in the data. The average d18O value for this period is 10.7 + 1.6‰ (n ¼ 25). The d18O values reach a maximum of –7‰ around 4 Ma. From 4 to 2 Ma, the d18O values are fairly uniform with an average of 8.8 + 1.2‰ (n ¼ 17).

Oxygen isotope ratio of pedogenic clay minerals Oxygen isotope ratios of pedogenic clay minerals have been used successfully to reconstruct past climate (Savin & Hsieh 1998). In Siwalik sediments, although smectite can be transported from the upper reaches of the Himalaya, higher abundances of this clay mineral in palaeosols compared to unaltered mudstone have been interpreted as evidence of their pedogenic origin (Stern et al. 1997). Pedogenic clay species such as smectite from the Mirpur and Bhaun sections of Potwar Plateau have been studied for their palaeoclimatic signatures (Stern et al. 1997). The d18O values of oxygen analysed from the aluminosilicate framework of smectite show two phases of lowering around 10 and 3 Ma, punctuated by an increased phase around 5.5 Ma (Fig. 4d). At around 10 Ma, the d18O value (VSMOW) is around 16‰ and subsequently these values show continuous enrichment (Fig. 6d). The maximum enrichment with a value of 19.6‰ was observed around 5.5 Ma. This enriched phase was followed by continuous lowering in d18O

Fig. 7. d18O values of diagenetic calcite cement from sandstone samples of Surai Khola. The two boundary curves one enveloping the maximum d18O values and the other the minimum d18O values are drawn. The first curve represents the effect of evolution of meteoric water. This shows an intensification of monsoon at around 6 to 7 Ma relative to the level 10 to 12 Ma. It also shows that monsoon system weakened at around 4 Ma and thereafter continued at the same level (up to 2.5 Ma). The d18O value of sandstone cement is in general lower compared to the d18O value of soil carbonate probably caused by precipitation of cement at higher temperature. The other curve represents evolution of fluid during burial. (Source: Sanyal et al. 2005a).

values, and at around 3 Ma, a value of 16‰ was observed again (Stern et al. 1997).

Late Quaternary continental records from the Indian sub-continent Fluvial sediments covering various parts of India have been used by various workers as important archives for deciphering monsoonal fluctuations during the Late Quaternary. Some of the most important regions from where such records are available include the Ganga plains, and alluvial plains of western India and Pakistan (Fig. 2). Furthermore, the Thar Desert (Fig. 2) forms another important region with respect to aeolian and lacustrine (playa) records covering the Late Quaternary and most importantly the post-glacial and Holocene periods.

Fan-terrace systems in intermontane basins Magnetostratigraphic data (Tandon et al. 1984; Cande & Kent 1995; Sangode et al. 1996; Kumarvel

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et al. 2005) suggest that the Early Quaternary fan sedimentation in the Himalayan foothills in the Chandigarh and Dehra Dun areas started around 1.77 Ma in a geological setting influenced by intraforeland thrusting (Kumar et al. 2007; Suresh et al. 2007). In areas such as Pinjor Duns in the northwestern Himalaya (Fig. 3), Quaternary fan and terrace deposits lie unconformably over Siwaliks and sedimentation in these areas for the last c. 100 ka has been influenced by both climate and tectonics (Suresh et al. 2007; Singh & Tandon 2007). Suresh et al. (2007) delineated two fan aggradation phases, Qf1 (c. 96– 84 ka) and Qf2 (c. 72 –25 ka) punctuated by an incision phase, which coincided with enhanced summer monsoonal precipitation around c. 74 ka (Thompson et al. 1997). Qf2 fan deposits also recorded a dominance of stream flow related facies, which correspond well with the known enhanced summer monsoon during the last interstadial (58–24 ka) in the Himalaya (Benn & Owen 1998). A retreat and toe cutting of Qf2 fan and associated deposits of axial river and lacustrine facies were attributed to local tectonics resulting in NE tilting. The Qf2 fan surface was incised between 20–16 ka during a weak summer monsoon, apparently in response to reduced sediment discharge and increased water/sediment ratio (increased stream power (Suresh et al. 2007). Post-Last Glacial Maximum (LGM) intensification of the summer monsoon resulted in deposition on terraces, T1 (upper, 16.5 ka) and T2 (lower, 6.5 ka), separated by a prolonged incision phase between 15 ka and 5 ka driven by Early Holocene monsoonal intensification. Singh & Tandon (2007) argued that tectonics rather than climate has played a significant role in controlling the incision in the Pinjor Dun area. The authors calculated a southwestern tilt on two of these fans, Kiratpur and Balad, producing uplifts of c. 41m and c. 46m at the fan heads for the Kiratpur and Balad fans, respectively. These figures broadly match tectonically-induced incisions of c. 45m measured at the fan heads. Late Quaternary alluvial fans in the Dehra Dun Re-entrant (Fig. 3) consist of four major stratigraphic units: the lower two units (c. 100 m thick) of mudflow and debris flow deposits with weak palaeosols overlain by multiple gravel beds of variable thickness (c. 30–100 m) deposited by braided streams and capped by more than 20 m thick overbank mudstones (Singh et al. 2001). The basal units of the fans composed of gravity flow deposits date between 40– 30 ka, and therefore, the time of initiation of these fans is taken as c. 50 ka. This has also been suggested to be the last major activity along the MBT/Bhauwala Thrust in this region. This phase continued until c. 10 ka and was then followed by deposition of braided stream deposits. It was terminated by the deposition of colluvial

deposits in response to a renewed uplift of the Dehra Dun block around 4 ka. The change over from gravity flow deposition to braided river deposition around c. 10 ka is related to Early Holocene monsoonal intensification (Singh et al. 2001). In a more recent work in the eastern part of Dehra Dun, Sinha et al. (2009) have reported terrace systems overlying Lesser Himalayan or Siwalik rocks. Geomorphological correlation, stratigraphic documentation and sedimentological analysis suggest four levels of terraces (T1 to T4) ranging in age from 11 ka and 0.5 ka. The upper two terraces (T3 and T4) terminate abruptly along the HFT and are present only in the hanging wall. The lower (T1 and T2) terraces cut across the MBT and HFT. These authors have interpreted a coupling of climate and tectonics in the development of these terraces. The incision events separating different levels of terraces occurred during 11 ka (T4 – T3), 9.7 ka (T3 –T2) and 6.9 ka (T2 –T1) out of which 11 ka and 6 ka events are consistent with widespread incision events in the Ganga plains (Gibling et al. 2005; Tandon et al. 2006; Sinha & Sarkar 2009) driven by monsoonal intensification. However, these authors have also implied a significant tectonic influence on these incision events due to movements along the HFT given the marked variation in their distribution across HFT.

Fluvial records from the southern Ganga plains The extensive Ganga plains provide by far the most important fluvial archive for reconstructing Late Quaternary monsoonal fluctuations. The stratigraphic information available for the upper 50 m of the alluvial cover studied in river bank sections (Sinha et al. 2005a, b; Gibling et al. 2005, 2008; Tandon et al. 2006) and drill cores (Sinha et al. 2007b; 2009) covering a time period of c. 100 ka reveals a complex history of dynamics of large and small river systems, and a strong climatic control. A marked geomorphological diversity from north to south, as well as east–west across the Ganga plains (Sinha et al. 2005c) adds further complexity in terms of the response of the river systems to external forcing. We constrain the fluvial history of the Ganga plains from several cliff sections in the southern Ganga plains and drill cores (see Fig. 8 for locations) described by earlier workers (Gibling et al. 2005a, b, 2008; Sinha et al. 2005, 2007b; Williams et al. 2006). Out of several cliff sections described by these workers, we choose two sections, which record a fluvial history of more than 100 ka. One of the cliff sections at Deoghat is 22 m high and located along the Belan River, a tributary

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Fig. 8. Major drainage lines in the western Ganga plains from where cliff sections and cores have provided the continental archives of Late Quaternary climate change.

of the Tons, which eventually meets the Ganga near Allahabad (Fig. 8). The second is the Kalpi section along the Yamuna River where 33 m of discontinuity-bound strata is exposed (Fig. 8) covering a time span of c. 120 ka (Gibling et al. 2005). These successions comprise a series of discontinuity-bounded sequences with alternate aggradational and degradational units. The aggradational units are characterized by overbank mud and sand, whereas the degradational units show gullied degradation surfaces, palaeosols, lacustrine mud, and aeolian silts. The Belan River has cut into the bedrock (Proterozoic quartzites) and the cliff section (Fig. 9) starts with channel-base calcretes above the Vindhyan Unconformity (Marine Isotopic Stage 5 or older) overlain by fluvial sediments (c. 85 to 16 ka, Marine Isotopic Stage 5 to 2). The fluvial record since c. 70 ka, consisting predominantly of muddy floodplain deposits with some meandering-river channel bodies, is in accord with generally high precipitation levels during Marine Isotopic Stage 3 to 5 (Gibling et al. 2008). Earlier work on the same section by Williams et al. (2006) suggested repeated aggradation and incision during this

period in response to variations in monsoon strength and sediment/water discharge. This period is also marked by widespread fluvial deposition elsewhere in the Ganga plains (Goodbred 2003; Srivastava et al. 2003b; Gibling et al. 2005). At the Kalpi section, Gibling et al. (2005) recorded aggradational sequences from c. 120 to 65 ka (Marine Isotopic Stage 5) and 90 to 40 ka (Marine Isotopic Stage 3 to 5) bound by major discontinuities and represented by carbonate-cemented sand and carbonate gravels. Deep gully fills within the floodplain sequence during 30 to 40 ka (Marine Isotopic Stage 3) indicate incision during this period. It was suggested that such ‘aggradationdegradation rhythms’ are in line with strong monsoonal fluctuations during this period; incision during wetter periods and aggradation during drier periods (Prell & Kutzbach 1987, 1992; Clemens & Prell 2003). The modern cliff line at Kalpi may have been created by incision during late Pliestocene/Holocene based on the correlation with other studies in the region (Williams & Clarke 1984; Goodbred & Kuehl 2000; Goodbred 2003); the available chronology did not permit this incision event to be constrained precisely.

164 P. SANYAL & R. SINHA Fig. 9. Representative lithologies of cliff sections and drill cores recording Late Quaternary monsoonal fluctuations in the Ganga Plains. (Source: Prell & Kutzbach 1987; Clemens & Prell 2003; Gibling et al. 2005, 2008; Sinha et al. 2007.)

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Both Kalpi and Belan sections have not preserved any record of the post-LGM (LGM)/ Holocene period apparently because of detachment and valley incision. However, a c. 12 m thick section from Mewar (MW in Fig. 8) along the Sengar River recorded two events of gullying around 12 ka and 10 ka, corresponding broadly with the Early Holocene monsoonal strengthening and consequently increased transport capacity of rivers (Gibling et al. 2005). Sinha & Sarkar (2009) have summarized the Late Pleistocene/Holocene response of the Ganga Basin and concluded that monsoonal fluctuations during the Holocene have played an important role in generating distinctive but variable alluvial architecture across the Ganga plains. In the Ganga Basin, valley margin sequences and valley fill cores in the Kanpur region (e.g. Firozpur core, Fig. 8) record two major channel sands corresponding to the pre-LGM and Holocene separated by intra-valley floodplain deposits (Sinha et al. 2007b). A multi-proxy approach using sedimentary facies, mineralogy and magnetic data for reconstruction of depositional environments during Late Quaternary suggests that valley aggradation occurred between 30 –15 ka during the climatic transition from a humid, pre-LGM to a cold-arid phase during the LGM. The adjoining cliff section on the valley margin (Bithur, see Fig. 8 for location) showed the accumulation of lacustrine and aeolian sequences around this time, indicating that floodplains were disconnected from the main channel around the LGM (Gibling et al. 2005). The Ganga River became ‘underfit’ (channel much smaller than the valley, Dury 1964) during this period with a much reduced channel capacity. Holocene monsoonal intensification rejuvenated the main flow of the Ganga River and a renewed phase of valley filling occurred during the transition from Early Holocene warm-humid conditions with higher precipitation and high sediment supply, to a relatively less humid period after 6 ka. Figure 9 presents summary columns with age dates from the Belan and Kalpi sections and representative cores from valley fill and interfluve settings. Two long proxy records for the Indian monsoon are also shown. The first is the upwelling and Indian summer monsoon intensity which is based on the stacking of five proxy records from cores in the Arabian Sea (Clemens & Prell 2003). The second is a record for modelled monsoon rainfall representing general trends for Asia (Prell & Kutzbach 1987). The cliff sections and cores from the Ganga plains broadly follow the modelled monsoonal curve; while aggradation in river valleys are noted during periods of strong summer monsoon, floodplain degradation and intra-valley floodplain accumulations mark the weak summer monsoon periods. Major discontinuities (periods of

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non-deposition during weak monsoon periods) are manifested as pedogenic events, gully fills, nonfluvial deposition or carbonate cementation/ calcrete development. Combining results from two representative sections at Belan and Kalpi, Marine Isotopic Stages (MIS) 5 (130 –70 ka) and 3 (60– 25 ka) are recognized as periods of major floodplain aggradation in the Ganga plains, which is consistent with the modelled monsoonal curve of high precipitation during this period. Cores from valley fills in the Ganga plains also suggest aggradational phases during the later part of Marine Isotope Stage 3. Marine Isotope Stage 2 is marked by a period of intra-valley floodplain development and pedogenesis in valley fills (Sinha et al. 2007b) and accumulation of aeolian and lacustrine sequences in valley margin sequences (Gibling et al. 2005). Most valleys in the western Ganga plains show that incision occurred in the late glacial/Early Holocene period in response to monsoonal intensification, which was primarily responsible for generating the modern landscape in this region. Several records from southern Asia (Sirocko et al. 1993; Overpeck et al. 1996; Srivastava et al. 2001; Pratt et al. 2002; Jain & Tandon 2003; Jain et al. 2005; Gibling et al. 2005, 2008; Williams et al. 2006; Sinha et al. 2007b; Clift et al. 2008) and in offshore areas (Goodbred & Kuehl 2000; Prins & Postma 2000; Goodbred 2003; Gupta et al. 2003; Chauhan et al. 2004; Rao et al. 2008) also suggest intensification of the monsoon following the LGM, with a peak in the Early to Mid Holocene, followed by a decrease in monsoon intensity. However, local variations in the precise timing of these events have been noted by several workers.

Fluvial records from western India Fluvial sequences in western India generated by the Luni, Sabarmati, Mahi and Narmada Rivers (Fig. 10) record a systematic variation of the sedimentation pattern during the last 100 ka in response to late Quaternary climate changes. Distinct phases of aggradation and incision are recognized in alluvial stratigraphy, and these events have been shown to be nearly synchronized in the lower reaches of the three river basins, the Luni, Mahi and Sabarmati Rivers (Jain et al. 2004). The oldest (MIS 5e) deposits in distal reaches of the Mahi River (Fig. 10) are marine clays, which are coeval with the muddy sequences in the inland reaches of the Sabarmati River (Fig. 10) and mark the wettest period in the region during the last 100 ka. The overlying gravels were interpreted as braided river deposits of c. 70 –75 ka age formed during Marine Isotope Stage (MIS) 5 to 4 transition (Juyal et al. 2000). The wet Marine Isotope Stage 3 is marked by two periods of prolonged fluvial

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Fig. 10. Major rivers in Western India and their climatic regime.

aggradation (52 –44 ka and 37 –30 ka) by mixed load meandering rivers and intermittent pedogenesis manifested as reddened horizons in both the Mahi (40 –25 ka, Juyal et al. 2000) and Sabarmati basins (58 –39 ka, Tandon et al. 1997). Little or no fluvial activity is recorded during late MIS 3 and 2, but aeolian dunes are prevalent in response to drier conditions during LGM. A re-establishment of monsoon is indicated by meandering river deposits during 14 –11 ka, followed by aeolian deposits corresponding to the drier phases in the last glacial maxima Late Glacial period. Both Mahi and Sabarmati Rivers were incised during the Early Holocene in response to monsoonal intensification. The Luni River sections (Fig. 10) show gravel bedload or gravel-sand bedload streams, which are inferred to have been active during the wet phase around 80 ka followed by periods of pedogenesis at 70 –30 ka. Progressive desiccation in the glacial period (30 –22 ka) led to the formation of ephemeral sand-bed rivers in this region (Jain & Tandon 2003; Jain et al. 2004); this period has been interpreted as a dominantly arid period with weak monsoonal conditions (Sirocko et al. 1993; Andrews et al. 1998). Although no definite signature of the LGM was recorded, one aeolian depositional phase was identified at c. 27 ka, which could mark the period of peak aridity.

Following the LGM, a renewal of fluvial activity corresponds to the strengthening of the summer monsoon. A diverse range of lithofacies representing gravelly braided streams, mixed load meandering streams, sand bed ephemeral streams, and aeolian process, dating at 14–11 ka are suggestive of a period of climatic instability. These deposits are inset with the older deposits (MIS 5 to 2), and therefore, an intervening period of incision was bracketed between 11 –9 ka. The period of incision is followed by sheet-flow deposition with some aeolian reworking between 9–5 ka and then an arid phase is recorded at 3 ka. Another incision event between 3 and 1 ka corresponds to a wet phase followed by a period of stability as indicated by slack water deposits (,1 ka; Kale et al. 2000) and accumulation of floodplain silts. The Narmada River in western India (Fig. 10) flows for a sizable part along the Narmada-Son Lineament, and the fluvial successions in this basin have been tectonically controlled throughout the Quaternary. However, a strong climatic signature has been recorded as sharp changes in fluvial style from multi-channel alluvial fan systems to meandering rivers in the late Pleistocene (Chamyal et al. 1997, 2002; Bhandari et al. 2005). The coastal systems of southern Gujarat show distinct effects of incision and aggradation forced by marine

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regression and transgression (Maurya et al. 2003), coupled with tectonic uplifts during the Late Holocene (Chamyal et al. 2003).

Aeolian records in Thar desert, northeastern India Climatic fluctuations during the Quaternary in the Thar Desert and the marginal areas preserved in aeolian records have been documented by several workers (Chawla et al. 1992; Thomas et al. 1999; Kar et al. 2001; Singhvi & Kar 2004) and supported by a robust luminescence chronology. The Early to Middle Pleistocene (.200 ka) period in Thar has generally been interpreted as relatively wet, followed by a distinct arid phase characterized by the development of extensive calcretes (Dhir et al. 1999, 2004), several playas, dunes and sand sheets. More intensive but episodic aeolian deposition, drying of river systems and lowering of groundwater level, has been recorded between 115 and 100 ka, suggesting an increase in aridity (Thomas et al. 1999; Singhvi & Kar 2004). Several other major phases of pre-LGM aeolian activity in the Thar Desert have been recorded at c. 75 ka, c. 55 ka and between 30 and 25 ka (Singhvi & Kar 2004). An interesting observation made by these authors is that at the LGM the Thar Desert experienced high aridity but wind strength was not sufficient for extensive aeolian accumulation. These authors further argued that the summer monsoon wind, which was responsible for dune building in the past, was very subdued during the LGM, which implies that the large scale aeolian dynamism did not coincide with LGM in the Thar, and therefore, resulted in dormancy of aeolian aggradation. Instead, a major period of aeolian deposition is recorded in Thar in the post-LGM period (13–11 ka) and has been identified as the Younger Dryas (YD) event in India, a period known for cooler climatic transitions and relatively higher winter rainfall. These interpretations are supported by studies on deep oceanic cores from the Arabian Sea (Sirocko et al. 1993; Overpeck et al. 1996; Gupta et al. 2003), which noted a lag of several years or centuries between peak rainfall and high wind speed. The Early Holocene monsoonal intensification also did not leave any distinctive signal in aeolian records in the Thar Desert, except for a period of ‘climatic optimum’ (7–6 ka) marked by stabilization of the sandy landscape and a weakly developed palaeosol layer (Singhvi & Kar 2004). The intervening period between Younger Dryas and the Climatic Optimum shows continued aeolian accumulation but with much less vigour. More recent phases of aeolian activity have occurred at 5.0–3.5 ka and 2.0–0.8 ka (Thomas et al. 1999) and a progressive northward

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shift in aeolian activity has also been recorded since the Early Holocene (Juyal et al. 2003). In the southern margin of the Thar Desert in Gujarat, aeolian sequences have recorded fluctuations in the summer monsoon since the LGM. Juyal et al. (2003) documented several dune profiles and the age data suggest two distinct episodes of dune building activities at 17–12 ka and 8–5 ka, separated by a period of landscape stability between 11 and 8 ka manifested as fluvial sands (reworked aeolian deposits). The former period of dune accumulation coincides with the time when the summer monsoon was strengthening in this region, thus confirming the earlier suggestions that dune accretion can occur during climatic transitions from cool and dry glacial period to wet Holocene (Wasson et al. 1983).

Calcretes in Thar Desert and Ganga Plains Calcretes, both soil and groundwater carbonate, are one of the striking landscape elements in Thar Desert and have been used as useful proxy for Late Quaternary monsoonal reconstruction. Andrews et al. (1998) studied the soil carbonate from the Thar Desert to reconstruct a detailed Late Pleistocene climate record using oxygen and carbon stable isotope data. The 4 m thick section near Shergarh in Thar Desert contained four aeolian units (EI– EIV) with abundant soil carbonate nodules and separated by colluvial units with sharp lower and upper contacts. The topmost aeolian unit (EV) did not contain any calcrete. The stable oxygen isotope data showed that Units EI (70 to 60 ka) and EIV (upper parts, c. 25 ka) are up to 4.4‰ higher than those measured for soil carbonate in Units E II (60 –55 ka) and EIII (55 –33 ka). Andrews et al. (1998) interpreted the aeolian Units I and IV as having formed during times of weaker monsoon and arid climate. Conversely, Units II and III were formed under strong monsoonal conditions. A strong covariance of oxygen and carbon isotope data was also noted. The carbon isotope data suggested that the local ecosystem was dominated by C4 vegetation (up to 88%) during times of weakened monsoon (70 –60 ka and c. 25 ka), whereas the period of enhanced monsoon (60– 25 ka) was marked by a reduction in C4 vegetation (75%). These observations are in line with the findings that ‘expansion of C4 grasses would be expected at low latitudes during glacial periods when atmospheric pCO2 was lowered’ (Cerling et al. 1997). Data published from younger calcretes (c. 19 to 2 ka) show more negative value of d13C ( –7.8 to –7.0‰) (Singhvi et al. 1996) suggesting that C4 vegetation was reduced even further to c. 28 –25% perhaps in response to strengthening of the summer monsoon after the LGM.

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Calcretes are abundant in Late Quaternary channel and floodplain strata of the southern Ganga plains. Srivastava (2001) interpreted three distinct climatic phases in the Late Pleistocene – Holocene based on the calcretes in the alluvial sequences in the Ganga-Ramganga interfluves and supported by clay mineralogical data (Srivastava et al. 1998). First, the period 12 920 to 7390 cal a bp was characterized by development of Type I carbonates characterized by irregular nodules of dense micrite and diffused needles. These carbonates are generally enriched in 12C (– 0.6 to –6.9‰), which is attributed to sparse vegetation and low rates of soil respiration. The host sediments contained trioctahedral vermiculite and smectite and septaric associated with illuvial clay features. These observations were used to interpret an arid climate between 13 and 7 ka in the Ganga plains. Secondly, the period 5730 to 4150 cal a bp was marked by Type II carbonates, which showed blocky calcite and needles in voids and enrichment in 13C (þ0.6 to þ1.8‰) because of extensive dissolution–re-precipitation (Srivastava 2001). These characteristics along with their association with smectite/kaolinite clay minerals are suggestive of a warm and humid climate. A third phase after 4150 cal a bp was manifested as secondary carbonate accumulation in voids interpreted to be groundwater related, and formed because of capillary rise in alternating moist and dry conditions with incomplete leaching (Srivastava 2001) accompanied by trioctahedral vermiculite and smectite in soils. This phase was also interpreted to reflect an arid to sub-humid climate in this region. Sinha et al. (2006a) described an unusually wide range of carbonate types associated with a Late Quaternary floodplain (interfluve) succession at Kalpi in the southern Ganga Plains (see Fig. 8 for location). At Kalpi, soil carbonates (nodules, rhizoconcretions, and powdery carbonate) are present within aggradational floodplain deposits, where they correspond to relatively high monsoonal precipitation and river discharge. In contrast, groundwater carbonates are associated with cemented degradational surfaces (discontinuities), which correspond with periods of relatively low precipitation. Mixed groundwater and soil carbonates are present in the deposits of small interfluve channels, and reworked nodules line degradational surfaces, and locally fill channels. From regional climatic modelling (Prell & Kutzbach 1987; Clemens & Prell 2003), the strength of the monsoon is known to have varied greatly during this period. The carbonates were used to test the response of the isotopic system and vegetation to precipitation changes over a period spanning much of Marine Isotope Stage 3 to 5. The d18O and d13C values of bulk and micro-drilled soil carbonate samples do not

show much variation through the sampled interval at Kalpi. Floodplain deposits were vegetated with a mixture of C4 and C3 plants (predominantly C4), with a higher proportion of C3 plants associated with channel deposits. This apparent lack of variation is surprising because the sampled interval represents at least 60 000 years of Marine Isotope Stage 3 to 5, when most of Asia experienced radical fluctuations in monsoon intensity and precipitation. Some of the apparent lack of variation was explained by preferential preservation of aggradational strata that represent relatively active monsoonal periods, as well as by the mixing of drier floodplain (C4) and wetter riparian (C3) vegetation. However, local departures from the regionally based climate model could not be ruled out. A modest up-section increase in C4 plants may represent increased aridity and lower atmospheric CO2. Isotopic analysis of organic matter from floodplain soil carbonate nodules (Sinha et al. 2006a) suggests a higher C3 plant contribution than carbonate-based data would suggest. The preserved organic matter may reflect the annual average biomass in the soil, whereas carbonate formation may have taken place mainly during the drier season when respiration of C4 plants was more important. It was interpreted that in interfluve settings such as Kalpi, seasonality may strongly affect the C3–C4 system, with preferential preservation of only part of the biomass.

Lacustrine records from Thar Desert and Ganga plains Lacustrine records for Late Quaternary climate change are available from two regions: (a) playas from Thar Desert in western Rajasthan that extend back to c. 30 ka and (b) from the Ganga plains which have provided high resolution records since 10 –15 ka. Most of the playas in the Thar Desert (Fig. 11) are filled with sediment and are usually dry at present. These have been studied by several workers to document climatic fluctuations during the Quaternary and the results are summarized in Figure 12. One of the first detailed studies (Singh et al. 1972, 1974) interpreted fluctuating climatic conditions in the Sambhar playa (Fig. 11) during the Holocene using several pollen species such as Typha angustata (a freshwater pollen), Syzgium Cuminii and Oldenlandia (both indicative of higher precipitation), Artemesia (characterizing .500 mm of rainfall) and Calligonum (grows in stabilized dunes). Singh et al. (1974) inferred five major climatic phases on the basis of these pollen assemblages: (a) arid (.11 ka); (b) arid with .250 mm rainfall (11–10.5 ka); (c) higher rainfall with east –west gradient (10.5 –6 ka);

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Fig. 11. Location map showing playas in the Thar Desert, western India which have provided important archives of Late Quaternary climate change.

(d) .500 mm rainfall in arid zone (6–3 ka); and (e) similar to present-day condition (,1 ka). Wasson et al. (1984) studied the Didwana playa (see Fig. 11 for location) employing mineralogical investigations of lake sediments. They interpreted hypersaline conditions around 13 ka, as indicated by a halite bed and low clastic sedimentation. A variable mineralogical assemblage between 13 ka and 6 ka reflects fluctuating climatic conditions (Fig. 12). The monsoon strengthened again around 6 ka, encouraging precipitation of dolomite and northupite and an increased clastic supply. The lake was dominantly fresh at this time and this condition continued until 4 ka. After this, the lake again started to dry up (Fig. 12). Enzel et al. (1999) studied lithostratigraphy and mineralogy of Lunkaransar playa (see Fig. 11 for location) to reconstruct a high-resolution Holocene environmental change and correlated the end of the wet period with a period of intense dune destabilization. Data indicate that a shallow fluctuating lake existed between 10 ka and 6 ka. After 6 ka there was a

sustained high-lake stand until 4.8 ka, as indicated by a few or no evaporite layers (absence of gypsum) in lake deposits and a decrease of d13Corg. The lake began to dry out after 4.8 ka (Fig. 12) and this coincides with an increase in aeolian sand coarser than 125 mm, a decrease in the CaCO3 percentage and the presence of broken, reworked calcite concretions. Sedimentological and mineralogical characteristics from the 6 m thick Bap-Malar playa (see Fig. 11 for location) sediments show a variation of climatic and hydrological variations during their deposition (Kajale & Deotare 1997). Weakly laminated bands of silt and clay at 4–6 m depth (.10 ka) indicate that the playa was ephemeral, and shallow, and that the occurrence of typical desert species pollen from these lower levels indicates relatively dry conditions during an early phase of sedimentation (Kajale & Deotare 1997). This was followed by a period of high salinity between 10 ka and 8 ka, as indicated by the presence of laminar gypsum with silt and clay. After 8 ka a lake high stand, with relatively

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Fig. 12. Palaeoclimatic history of Thar desert based on lacustrine records (data source: Wasson et al. 1984; Kajale & Deotare 1997; Enzel et al. 1999; Kale et al. 2003; Sinha et al. 2006b).

freshwater conditions is indicated by welllaminated silt and an absence of gypsum. This condition continued until about 5 ka, after which the lake started to dry out (Fig. 12). Sinha et al. (2006b) investigated a c. 23 m deep borehole from the Sambhar playa and used an integrated mineralogical and geochemical approach to derive information regarding palaeohydrological and palaeoclimatic fluctuations in the Thar Desert for the last c. 30 000 years (Fig. 12). The Sambhar Lake came into existence well before the LGM and much earlier than the other playas in the Thar Desert. The oldest (.30 ka) lacustrine sediments overlying weathered calcrete deposits indicate a shallow lake with low to moderate salinities during the period c. 30 to 24 ka. The available records from western Rajasthan from calcretes (Andrews et al. 1998), alluvial plains (Jain & Tandon 2003) and aeolian dynamism (Chawla

et al. 1992) also suggest less arid conditions and a normal summer monsoon with reference to modern conditions. Lacustrine sediments at c. 12 –10 m depth (c. 24 –20 ka) characterize a relatively arid phase, which corresponds to the weakening of the summer monsoon during the LGM. A sharp increase in d18O values and [MgO/ (MgO þ CaO)] ratio and moderate values of Na/ Al recorded at c. 10 m depth reflect low chemical weathering, and therefore, suggest an increase in aridity. Based on these data, it was suggested that Sambhar was not desiccated during the LGM and continued to hold some water perhaps because of its location at the desert margin (Sinha et al. 2006b). In the postglacial period (c. 20–16 ka), a shallow lake condition (Fig. 12) is indicated by lower Na/Al ratio, depleted d18O values and enriched d13C values. This observation is in line with the modelling results (Prell & Kutzbach

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1987) suggesting gradual strengthening of the Indian summer monsoon after 20 ka reaching a maximum around 10 ka. The period c. 16.0 to 7.5 ka is marked by variable salinity, higher chemical weathering and higher organic productivity, and therefore, suggests semi-humid conditions. This period is close to the ‘largest event’ of Sirocko et al. (1993) at 16.06 ka, which was correlated to the beginning of global Melt Water Pulse Ia. This phase also includes the abrupt monsoonal intensification in the region in several steps, four steps between 17.7 and 9.9 ka (Sirocko et al. 1993) and two steps between 13 and 9.5 ka (Overpeck et al. 1996). At Sambhar, a rise in the lake level is recorded between 6.8 and 2.5 ka (Fig. 12) as evidenced by gradual depletion of d18O values and return of calcite-rich facies. This period broadly matches the peak monsoon predictions during c. 9.5–5.5 ka by most workers and also with the high stand in Lunkaransar (c. 7.2 to 5.6 ka, Enzel et al. 1999) and Didwana (c. 6–4 ka, Wasson et al. 1984). The lake level at Sambhar was lowered again after 2.5 ka due to the effects of Middle Holocene increase in aridity and gradual weakening of summer monsoon rains. This is manifested in higher amounts of sulphate minerals such as polyhalite (K2Ca2Mg (SO4)4 . 2H2O) and thenardite (Na2SO4), although calcite and halite still dominated the assemblage. The onset of this aridity at Sambhar lags by more than 1500 –2500 years in comparison to that seen at Didwana (c. 4 ka, Wasson et al. 1984) and Lunkaransar (c. 5 ka, Enzel et al. 1999) playas, where water level started to drop much earlier. Furthermore, both Didwana and Lunkaransar playas were completely desiccated at 3– 4 ka. In contrast, the Sambhar record does not record any evidence of complete desiccation and has remained ephemeral until the present. Such spatial differences are attributed to the location of the Sambhar playa at the desert margin. Not only is the present-day mean precipitation higher at Sambhar compared to Didwana or Lunkaransar playas, but the simulation models (Bryson & Swain 1981; Swain et al 1983) for the Thar region also predict such spatial variations. Limited lacustrine records for the Holocene period are available from the Ganga plains. In the Sanai Tal area in the central Ganga plains located c. 50 km NE of Kanpur (see Fig. 8), Sharma et al. (2004a, b) observed a dry spell between 5 and 2 ka, in line with regional-scale climate models from the Thar Desert in western India, which reported lake desiccation at about 4.8 ka (Swain et al. 1983; Singh et al. 1990) and enhanced dune building activity at about 5 ka (Thomas et al. 1999). Furthermore, the Holocene stratigraphy in the alluvial plains often shows intercalations of

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aeolian and lacustrine deposits (Srivastava et al. 2003b), which indicate high-frequency changes in the hydrological system, apparently triggered by monsoonal fluctuations.

Discussion Monsoonal reconstructions from oxygen isotope ratio of soil carbonate and pedogenic clay minerals from Siwalik successions The oxygen isotope ratio of soil carbonate and pedogenic clay minerals depends on the temperature of their formation, and isotopic ratio of the soil water which in turn is a function of local rain water composition. The temperature of carbonate precipitation and formation of pedogenic clay minerals can be taken as representative of mean annual temperature of an area. In the past, the climate was characterized by repeated fluctuations in temperature. In the recent geological past, the maximum lowering occurred during the LGM (c. 18 –23 ka). The d18O variability obtained from speleothem data suggests that during the LGM temperature in the tropical region was lowered by 3– 4 8C (Farrera et al. 1999). This temperature change can impart c. 0.8‰ enrichment in the d18O value of soil carbonate and smectite. As the observed variations in d18O value are more than 0.8‰, it can be assumed that the d18O values are mainly controlled by the d18O value of rain water, which implies that temperature has not played a major role in determining variations of oxygen isotope ratio in Siwalik samples. Hence, the d18O value of soil carbonate and smectite can be used to decipher the average isotopic composition of rain water (Quade et al. 1989). The d18O value variations at different Siwalik sections suggest three phases of lowering at around 10.5, at 5.5 and at 3 Ma, which correspond to periods of intensification of the Indian summer monsoon (see Fig. 6a–d based on Quade et al. 1989, 1995; Kroon et al. 1991; Stern et al. 1997; Sanyal et al. 2004, 2005b). The summer monsoon in the Indian subcontinent is currently characterized by heavy rains during June, July, August and September and is caused by convective vortex of WNW moving depressions originating from the Bay of Bengal operating on moist oceanic air moving in the same direction (Fig. 1). In contrast, during the winter, dry continental air blows from the northeast, resulting in little rain except occasional precipitation in the Siwalik region due to western disturbances bringing moisture from the Mediterranean region (Webster et al. 1998). The winter rains are usually more enriched in 18O compared to summer rains, as observed from

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the seasonal variation data of Delhi rains (IAEA 2003). Furthermore, d18O values of rainwater decrease with an increase in the amount of precipitation (IAEA 2003). At low latitudes, the average monthly rainfall and the mean monthly d18O values are usually negatively correlated, where an increase in 100 mm of precipitation is associated with a lowering in d18O value by 1.5‰ (Yurtsever & Gat 1981). Therefore, it can be argued that the phases of depletion in the oxygen isotopic composition of carbonate and clay minerals were the result of contemporary intensification of summer monsoon because intensified monsoonal wind system would generate more intense and frequent depressions (storms) and result in more rains.

maximum observed d18O value at any level (or age) represents calcite precipitation near the surface. In Figure 7, two curves are shown, one enveloping the maximum d18O values and the other the minimum d18O values. The first curve should represent the effect of evolution of meteoric water with the lowest values observed around 6–7 Ma (intensification of monsoon) compared with the time period 10–12 Ma. However, the most enriched values are observed in the time range 4 –2 Ma indicating a relatively drier period. The other curve of the enveloping surface represents evolution of fluid during burial (Fig. 7).

Oxygen isotope ratio of carbonate cement in sandstones from Siwalik successions

Monsoonal rainfall variations reconstructed from the oxygen isotopic ratios of soil carbonate, smectite and carbonate cement of sandstone from the Siwalik successions are also reflected in other proxies such as pollens, geochemical parameters and magnetic mineralogy. Variations in the different taxa documented from the palynological study from the Surai Khola section of Nepal showed change in the vegetational pattern in response to monsoonal change since 11.5 Ma (Hoorn et al. 2000). During 11.5 –8.0 Ma, dominance of broad-leafed taxa indicates subtropical to tropical and temperate climates. Replacement of subtropical and temperate forests by grasslands between 8.0 and 6.5 Ma reflect more seasonal drought. Although, an increase in sediment discharge in the fluvial system is inferred from an increase of the aquatic taxon Potamogeton at 8–6 Ma. Evidence of climatic cooling between 6.5 and 5 Ma is indicated by the appearance of steppe taxa, lowering of tropical forest taxa and disappearance of certain Dipterocarpaceae. From 5.5 to 3.3 Ma an abundance of algae (Spirogyra) and pteridophytes in the pollen assemblages indicate increased seasonal flooding which may have produced local lacustrine conditions on the overbanks. Overall, pollen data also indicate two phases of monsoon intensification around 10.5 and 5.5 Ma punctuated by a dry phase around 6 Ma. Additionally, Himalayan foreland basin sediments record the appearance of C4 plants over a time period of 9 to 6 Ma (Quade et al. 1989, 1995; Sanyal et al. 2004, 2005a) which is closely linked to the monsoon dynamic in the Indian subcontinent. The pollen record from the Haripur Khol section since 4 Ma shows variation in the growing season conditions. From 4.0 to 3.5 Ma the dry grassland indicates an arid climate. The presence of Amaranth/ Chenopodiaceae along with members of Polypodiaceae and other ferns further shows that in a few places, these plants were thriving during rainy to

Compositionally, Siwalik sandstone is typically lithic arenite with calcite as a cementing material. Calcite is precipitated in the sub-aerial vadose zone due to capillary action under wet and dry climatic conditions (Tandon & Varshney 1991) and occurs in a pore-filling variety with a non-uniform distribution in sandstone. The d18O value of calcite cement in sandstone depends on the temperature of its crystallization and on the d18O value of the ambient water. Calcium carbonate precipitating at shallow levels should reflect the d18O value of the local groundwater, which is normally derived from local rainfall. However, subsequent burial of sandstone can alter the original value. Therefore, if the d18O value contribution due to burial can be isolated then the d18O value of carbonate cement associated with the sandstone can be used to determine the temporal rainfall variations. Because increase in temperature lowers the d18O value, the lowest value in a suite of samples may correspond to the highest temperature experienced by the carbonate during burial. A large scatter in d18O values (from –10 to –18‰) at the bottom (12–10 Ma) of the stratigraphic succession may be due to dissolution and re-precipitation of calcites at different temperatures corresponding to various depths and generations during burial. Cementation at different times of burial could also cause the scatter. During the burial process the d18O value of cement could have changed because of dissolution/ re-precipitation of calcite. However, some samples experienced minimal change, with some of them even retaining the original d18O values. This implies that the lowest d18O value is the most altered and the highest value is the least altered when samples are compared at the same stratigraphic level. For simplicity, it was assumed that the

Other evidences of monsoonal fluctuations from Siwalik successions

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early summer times. From 3.5–2.7 Ma and the latter half of this stage, significant rise in Amaranth/Chenopodiaceae and large concentration of Ceratopteris pores indicates a marshy or muddy condition indicating increase in rainfall. During 2.7–2.5 Ma, the spores of Lycopodium, algal cysts as well as dicot pollens suggest the existence of well-developed ponding conditions. Presence of Nymphaeaceae indicates shallow-water or seasonal lacustrine habitat. Pinus pollen also indicates a temperate to subtropical climate. During 2.5–1.0 Ma, the observation of abnormally thick cell walls is intriguing but could be due to a response of the leaf to a cool and dry climate (Phadtare et al. 1994). The above discussion indicates that the post-4 Ma vegetational regime in this part of the Indian Siwalik shows signature of monsoonal changes. Climatic variations are also observed in the various chemical and physical properties of the Haripur Khol palaeosols. Climatic fluctuations have been inferred (Sangode et al. 2001) from the variation of oxidation, hydroxylation and humification index of palaeosols as measured from magnetic mineralogy. The indices are based on selective saturation levels of induced magnetic field, inorganic and organic carbon content, and Rb/Sr ratios. Variations in the indices indicate large-scale climate changes within the Pliocene–Pleistocene, with a warm-humid climate during the Early Pliocene and intermediate phase during the early MidPliocene and warm oxidative phase during the Mid to Late Pliocene. The early –mid Pleistocene was characterized by a cold phase (Sangode et al. 2001).

Monsoonal response to fluvial architecture in Siwalik successions Atmospheric general circulation models show that the uplift of the Tibetan Plateau was one of the most crucial factors in initiating the monsoon in the Indian subcontinent (Prell & Kutzbach 1992) and subsequent studies have also demonstrated that phased uplift of the Himalayan-Tibetan Plateau played a major role in the evolution of Indian monsoon (Zhisheng et al. 2001). Compilation of data on exhumation and erosional chronology (White et al. 2001; Harris 2006; Yin 2006 and references therein) of the Himalaya shows major phases at 25–10 Ma, 6–5 Ma and 3 –2 Ma throughout the western, central and eastern Himalayas. These phases are characterized by activation/reactivation of Main Central Thrust, Main Boundary Thrust, Great Counter Thrust, and South Tibetan Detachment System (Fig. 2). Tectonic activity in these major structural elements is also reflected in the increase in sedimentation rate, amalgamation of sand bodies, gravel progradation and contribution

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of Greater Himalayan crystalline minerals/lithic fragments in the foreland basin as well as on the Bengal Fan. Kumar et al. (2003a) observed two major changes of sedimentation pattern and drainage organization at 10 Ma and 5 Ma and several interspersed events during the 10–0.5 Ma period. The succession around 10 Ma is characterized by multi-storied sandstone with abundant erosional surfaces but no lateral accretionary surfaces, and low palaeoflow variability. These features indicate that deposition took place in a regularly avulsing large braided river system. The net sediment accumulation rate also increased by a factor of 2 to 3. Mass accumulation rate also shows rapid increase around 10 Ma in the Ganga Basin (Metivier et al. 1999) and Indus Fan (Clift 2006). Burbank et al. (1996) showed that almost all the Siwalik sections from Pakistan to Nepal recorded acceleration in sedimentation and increase in subsidence at around 11 Ma. This increase in sediment flux in the Himalayan foreland was probably due to combined action of intensified monsoonal precipitation and tectonic activity. On the other hand, the presence of halite in the Ranital section (Fig. 2) around 6.5 Ma indicates extreme aridity, which marks a weak phase of monsoon. Subsequently, widespread distribution and stacking of conglomerate and progradation of alluvial fans was observed between 6 and 5 Ma. Studies carried out by Retallack (1995) in a few other Siwalik sections of Pakistan and India also showed variations in lithofacies with time. Thick amalgamated sandstone with minor overbank deposits without humic horizons prior to 8 Ma and abundant overbank deposits with shallowly leached palaeosols with significant humic horizons in post-8 Ma sediments in Pakistan and the Indian Siwalik sections suggest that flooding and effective rainfall was considerably greater before 8 Ma (Retallack 1995). Additionally, an iron concretion layer within calcareous palaeosols in sediments older than 8 Ma probably represents wet season whereas calcareous layers may indicate dry season deposition. The presence of unusually diffused calcic horizons reflects dry season variation in the wetting front within the soil profile (Retallack 1995) typical of a monsoonal regime. Elsewhere, a higher monsoonal activity before 8 Ma is also evident from higher chemical weathering in Bengal Fan (Derry & France-Lanord 1996) and in the Indus and South China Sea (Clift et al. 2008).

Spatial variability of monsoon as manifested in Siwalik records Variations of oxygen isotope ratio in soil carbonate, pedogenic clay minerals and carbonate cement

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indicate fluctuations in the intensity of monsoonal rainfall since 18 Ma. Although the intensification of monsoon around c. 10.5 Ma is observed in all sections, the other two phases of intensified monsoon at 5.5 and 3 Ma are recorded preferentially in some sections. After 6 Ma, no coherent trend is observed in Potwar Plateau because of high scatter in data. In case of Indian Siwalik, d18O values showed a clear swing towards less negative values after 5.5 Ma suggesting a shift of the climatic condition towards a wetter phase This sharp change is probably present but not clearly recorded in the profile of Quade et al. (1989) (Fig. 6a). In case of the Surai Khola section of the Nepalese Siwalik, as mentioned above, subsequent to 6 Ma, the d18O value of carbonate cement of sandstone increases rapidly, which is consistent with the observed enrichment of 18O of soil carbonate in the Surai Khola section (Fig. 6b; see also Quade et al. 1995) and other sections of the Indian Siwalik (Sanyal et al. 2004). The profile of lowest d18O values, which corresponds to the evolution of meteoric water seems to agree with the isotope data of soil carbonates, although the latter set has a large scatter. Study of soil carbonate from other parts of the Indian subcontinent also supports intensification of monsoon around 5.5 Ma (Sanyal et al. 2004). This indicates that despite alteration and consequent scatter, a signature of rapid d18O change can still be retrieved from sandstone data. Additionally, it establishes that the d18O signal is definitely induced by meteoric water and represents a significant regional change in the rainfall pattern. The d18O value of sandstone cement is generally lower than that of soil carbonate at the same stratigraphic depth, though a similar trend in their depth variation is observed. As mentioned above, this can be explained in terms of higher temperature of formation for the sandstone calcite. It is possible that reduced d18O values in sandstone cement can also be caused by mixing of river water with shallow groundwater by lateral flow. In the Himalayan region, river water flowing from the highlands has lower d18O values than the water from lowlands. Mixing between groundwater and the local river water would therefore make the d18O value of formation water more depleted.

fluctuations in river systems (Gibling et al. 2005), and that the balance between discharge and sediment load would have governed the energy available for erosion. Depending on this balance, the rivers aggraded/degraded their floodplains by remaining connected or disconnected from their floodplains. Such large-scale connectivity of floodplains with the main river (Fig. 13) is primarily linked to increased or decreased stream capacity in the major rivers in response to monsoon precipitation (and discharge) supported by chronological evidence (Gibling et al. 2005). For example, during a decline in monsoonal precipitation from Marine Isotopic Stage 3 into the LGM in Marine Isotopic Stage 2, the parent river would inundate its floodplain much less frequently and would possibly become ‘underfit’ within its valley and disconnected from its floodplain (Fig. 13). This may be manifested as a change in facies from floodplain

Fluvial response to Late Quaternary monsoonal fluctuations: forcing functions and stratigraphic manifestations

Fig. 13. A model for floodplain aggradation and degradataion from the main river as a function of monsoonal fluctuations: (a) floodplains are connected to the main during high monsoonal phase and aggrade; (b) floodplains are disconnected either due to underfitting during low monsoon or incision during high monsoon and degradation takes place in various forms; and (c) composite stratigraphy of floodplains evolves through such aggradation-degradation rhythms. (Redrawn from Gibling et al. 2005).

In the Ganga plains, the response of the river systems to monsoonal fluctuations since 20 ka has been recorded as aggradation-degradation rhythms in the floodplains. It has been argued that monsoonal fluctuations were accompanied by strong discharge

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to non-fluvial deposits (aeolian, lacustrine) as recorded at Bithur section along the Ganga River (Fig. 8). An alternate scenario is that the river would have incised during a period of increasing precipitation for example, late MIS 3 (30 –40 ka) or in the Early Holocene, and probable increased transport capacity of rivers (Fig. 13). In addition to promoting floodplain reworking, increased stream capacity would caused incision of the main river and its tributaries leading to gully erosion and badland formation in upstream reaches. Such incision events relating to increased monsoonal precipitation have been recorded at the Kalpi section along the Yamuna River (c. 30 ka, Fig. 8) and at the Mewar section along the Sengar River (c. 9 ka, Fig. 8). During long periods of disconnectivity, some parts of the floodplain may still continue to aggrade through reworking by plains-fed rivers or through occasional flooding. A recent study by Rahman et al. (2009) has demonstrated that such monsoon-driven changes in the hydrological regime of the hinterland generates a signal in terms of shifts in sediment provenance as far as distal Ganga plains. Both 87Sr/86Sr and 1Nd values for the IITK core (see Fig. 8 for location) showed major incursions at 70 ka and 20 ka coinciding with lower monsoon intensities and maximum glacial cover, thereby limiting the sediment supply from the Higher Himalaya (Rahman et al. 2009). The fluvial successions in western India have also recorded a systematic variation of the sedimentation pattern in response to Late Quaternary climate changes. The phases of aggradation and incision occurred around the same times in the lower reaches of the three river basins, the Luni, Mahi and Sabarmati. Variations in fluvial styles in different regions are apparently a function of the precipitation gradient, which is observed even today.

Fluvial response to Late Quaternary monsoonal fluctuations: spatial heterogeneity and controlling factors Available data from different parts of the Indian sub-continent demonstrate a strong climatic influence on the stratigraphic development during the Late Quaternary (Fig. 14). In the Ganga plains, significant spatial heterogeneity across the Ganga plains as a function of precipitation variability, along-strike variability in hinterland tectonics, and sediment flux suggest ‘differential sensitivity’ of river systems to monsoonal fluctuations (Sinha et al. 2005c). As a result, we record a complex landscape response of the river systems in the Ganga Plains along an east –west transect. For example, the western parts of the Ganga Plains developed incised valleys during the Early Holocene that are

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partially filled while the eastern Ganga Plains are presently characterized by a regime of rapid aggradation (Sinha et al. 2005c) and high floodplain sedimentation rates (c. 1.5 mm a21; Sinha et al. 1996). Presuming that such geomorphic diversity existed for most of the Late Quaternary, individual components of large river systems such as the Ganga would have responded differently to monsoonal forcing. Data on the stratigraphic response of the Gangetic rivers to monsoonal forcing presented in the previous sections have demonstrated that the stable configuration of valleys and interfluves in the western plains has generated discontinuitybounded sequences (Fig. 13) in wide interfluves (Gibling et al. 2005) and multiple stages of valley filling episodes in narrow, incised valleys, broadly corresponding with climatic transitions (Sinha et al. 2007b) during the Late Quaternary. The manifestation of the discontinuities in different exposed sections is varied and depends upon the geomorphic setting. For example, in valley margin settings such as Bithur (Fig. 8), it is manifested as non-fluvial deposition (aeolian, lacustrine) whereas in proximal to distal interfluves settings such as Kalpi and IITK (Figs 8 & 9), it is manifested as palaeosol development or calcrete formation. On the other hand, the eastern Ganga plains in north Bihar, characterized by large fans and interfan areas, have generated large, multi-storied sand sheets below the fans and near-continuous, thick overbank muds below the interfan areas (Jain & Sinha 2003b; Sinha et al. 2005b). Another example of spatial inhomogeneity of river response to monsoonal fluctuations is noted between desert and desert-margin rivers in western India. For example, the Luni River was very dynamic during the Late Glacial (14–11 ka) period and experienced frequent transitions from gravel bedload braided streams to ephemeral sand-bed rivers and suspended load-dominated meandering streams. The coeval deposits in the Sabarmati River, however, suggest a meandering, floodplaindominated river during this period. The desert streams thus appear to be more sensitive to climate change, and may have very short response times and low geomorphological thresholds as compared to the desert-margin rivers (Jain & Tandon 2003; Jain et al. 2004).

Holocene monsoonal variability and responses High-resolution records available from the Holocene period suggest that the Southwest Monsoon peaked at 10 –5.5 ka (Overpeck et al. 1996; Fleitman et al. 2003), in accord with pollen records

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Fig. 14. Summarized view of the response of terrestrial systems to Late Quaternary monsoonal fluctuations (compiled from various sources mentioned in the text).

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from Tibet (van Campo & Gasse 1993). In contrast, several records from India suggest that the monsoon intensification peaked after 8 –9 ka and perhaps at 6 ka or even later. In Ganga plains, the valley cores such as Firozpur (Figs 8 & 9) have not preserved any signature of Late Holocene aggradation (Sinha et al. 2007b). One apparent reason is the intra-valley migration of the river and incision which is evident in the modern landscape (Sinha et al. 2007b). However, our recent unpublished results from cores raised from reaches 35 –40 km upstream of Kanpur/Bithur show that these valleys aggraded very rapidly during the Late Holocene generating c. 8 to 10 m of channel sands in a period of ,3 ka (N.G. Roy, pers. comm. 2009). This is possibly an indicator of a sharp decrease in transport capacity of rivers in response to recent decline in monsoon intensity coupled with anthropogenic activities, which may also have increased the sediment supply. In contrast, the eastern parts of the Ganga plains in northern Bihar (Fig. 2) and the fluvio-deltaic plains in the Bengal Basin (Fig. 2) show extensive and rapid aggradation for most of the mid to Late Holocene period (Sinha et al. 1996; Sinha & Sarkar 2009) and a very dynamic fluvial regime (Sinha 1996; Jain & Sinha 2003a, 2004), perhaps a manifestation of low accommodation space in an aggrading regime. Chandra et al. (2007) also recorded rapid aggradation and incision in the Ghaghra-Rapti interfluve (Fig. 2) between 11.5 and 5.5 ka in an avulsive regime. In the Bengal Basin, an increase in sediment supply in the Bay of Bengal (2.3 times more than the modern flux) between 11 and 7 ka has been correlated with Early Holocene monsoonal strengthening (Goodbred & Kuehl 2000). Several records from the Lower Ganga plains and delta plains have documented repeated aggradational phases and enhanced chemical weathering under warm and humid conditions in response to climatic fluctuations, in addition to the effect of seal level changes (Goodbred 2003; Heroy et al. 2003; Chauhan et al. 2004; Sinha & Sarkar 2009). Evidence from Thar Desert playas suggest that precipitation increased to a maximum at c. 7.2 –6.0 ka (Prasad & Enzel 2006), and Himalayan peat records suggest that the monsoon strengthened after 7.8 ka to a peak between 6.0 and 4.5 ka (Phadtare 2000). As a result of orographic precipitation on the Western Ghats, Peninsular India experiences reduced rainfall and the monsoon may have intensified in this region only after 8–9 ka (Thamban et al. 2001; Staubwasser & Weiss 2006). As noted by Prasad & Enzel (2006), Asia receives precipitation from both the summer and winter monsoon, and lake and river systems are affected by local temperature and evaporation,

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groundwater supply, and regional precipitation gradients. Additional climatic perturbations arise from mid-latitude westerlies and El Nin˜o Southern Oscillation (Owen et al. 2005; Williams et al. 2006), and local topography and rain-shadow effects may influence thresholds for response recorded in the proxy records.

Conclusions The uplift history of the Himalaya in general and evolution of the Tibetan Plateau in particular, have played a critical role in the evolution of the Indian summer monsoon. The Himalayan orogeny and the subsequent growth of several fault systems have controlled the fluvial sedimentation in the Himalayan foreland in a major way. The Siwalik successions exposed along the strike of the Himalaya therefore provide the longest terrestrial records of monsoonal fluctuations. Although Siwalik successions in the foreland continue through the Quaternary period, no climatic reconstructions have been possible due to the lack of any proxy in gravels and sands of the Upper Siwaliks. In distal parts of the foreland basin in the Ganga plains, the Quaternary successions are very thick (.4 km in certain parts). Lack of good exposed sections in this region has always precluded the development of long-term stratigraphic framework and palaeoclimatic interpretations. In northwestern and western India, fluvial and aeolian records have unravelled monsoonal fluctuations for the last c. 120 ka but the most impressive records are available from the Thar Desert playas for the last 30 ka. The most striking feature of late Quaternary records from the Indian sub-continent is a marked spatial variability in terms of stratigraphic manifestation, which is a function of geomorphological diversity and rainfall variability which exists even today. Future work in this region must focus on penetrating deeper stratigraphy and generating a sound chronological framework. The authors profusely thank Peter Clift for the invitation to write this paper. The synthesis presented in this paper is an outcome of several collaborations involving people from different institutions: S. K. Tandon (Delhi), Martin Gibling (Dalhousie), Mayank Jain (Riso), S. K. Bhattacharya (PRL, Ahmedabad), Rohtash Kumar, S. K. Ghosh all from Wadia Institute of Himalayan Geology, Dehra Dun), Mahesh Prasad (BSIP, Lucknow), S. J. Sangode (Nagpur) and funding from several agencies such as Department of Science and Technology (New Delhi), NERC (Canada), and IIT Kanpur. We thank all our collaborators for the intense discussions and funding agencies for their support. RS thanks his students, Ananda Shanker Dasgupta, N. G. Roy, Parthasarthi Bhattacharjee, G. Babu and several others who contributed through their thesis work.

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Monsoon control over erosion patterns in the Western Himalaya: possible feed-back into the tectonic evolution PETER D. CLIFT1*, LIVIU GIOSAN2, ANDREW CARTER3, EDUARDO GARZANTI4, VALIER GALY2, ALI R. TABREZ5, MALCOLM PRINGLE6, IAN H. CAMPBELL7, CHRISTIAN FRANCE-LANORD8, JUREK BLUSZTAJN2, CHARLOTTE ALLEN6, ¨ CKGE9, MOHAMMED DANISH5 & M.M. RABBANI5 ANWAR ALIZAI1, ANDREAS LU 1

School of Geosciences, University of Aberdeen, Aberdeen, AB24 3UE, UK 2

Woods Hole Oceanographic Institution, Woods Hole, MA 02543, USA

3

School of Earth Sciences, University and Birkbeck College London, Gower Street, London, WC1E 6BT, UK

4

Dipartimento Scienze Geologiche e Geotecnologie, Universita’ di Milano-Bicocca, Piazza della Scienza 4 – 20126 Milano, Italy 5

National Institute for Oceanography, Clifton, Karachi 75600, Pakistan

6

Department of Earth, Atmospheric and Planetary Sciences, Massachusetts Institute of Technology, Cambridge, Massachusetts, USA 7

Research School of Earth Sciences, The Australian National University, Canberra, A.C.T. 0200, Australia

8

CRPG-CNRS, BP 20, 15 rue Notre Dame des Pauvres, 54501 Vandoeuvre les Nancy, France 9

Bundesanstalt fu¨r Geowissenschaften und Rohstoffe (BGR), Stilleweg 2, D-30655 Hannover, Germany *Corresponding author (email: [email protected])

Abstract: The Indus Delta is constructed of sediment eroded from the western Himalaya and since 20 ka has been subjected to strong variations in monsoon intensity. Provenance changes rapidly at 12–8 ka, although bulk and heavy mineral content remains relatively unchanged. Bulk sediment analyses shows more negative 1Nd and higher 87Sr/86Sr values, peaking around 8 –9 ka. Apatite fission track ages and biotite Ar–Ar ages show younger grains ages at 8– 9 ka compared to at the Last Glacial Maximum (LGM). At the same time d13C climbs from – 23 to – 20‰, suggestive of a shift from terrestrial to more marine organic carbon as Early Holocene sea level rose. U–Pb zircon ages suggest enhanced erosion of the Lesser Himalaya and a relative reduction in erosion from the Transhimalaya and Karakoram since the LGM. The shift in erosion to the south correlates with those regions now affected by the heaviest summer monsoon rains. The focused erosion along the southern edge of Tibet required by current tectonic models for the Greater Himalaya would be impossible to achieve without a strong summer monsoon. Our work supports the idea that although long-term monsoon strengthening is caused by uplift of the Tibetan Plateau, monsoon-driven erosion controls Himalayan tectonic evolution. Supplementary material: A table of the population breakdown for zircons in sands and the predicted Nd isotope composition of sediments based on the zircons compared to the measured whole rock value is available at http://www.geolsoc.org.uk/SUP18412

The relationships between climate, continental erosion and mountain building continue to be debated and are central to understanding how the solid planet and its atmosphere have interacted over long periods of geological time. In particular, the

links between mountain building in Cenozoic Asia and the intensification of the monsoon are presently unclear, with the different processes feeding back on each other. Although climate modelling has shown that a wide, high Tibetan Plateau is important to

From: Clift, P. D., Tada, R. & Zheng, H. (eds) Monsoon Evolution and Tectonics –Climate Linkage in Asia. Geological Society, London, Special Publications, 342, 185–218. DOI: 10.1144/SP342.12 0305-8719/10/$15.00 # The Geological Society of London 2010.

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maintaining a strong monsoon circulation (An et al. 2001; Kitoh 2004) it is also clear that monsoon strength has varied greatly over millennial to orbital timescales, driven by solar factors either directly or via its influence on the intensity of northern hemispheric glaciation (Clemens et al. 1991; Wang et al. 2005). Feedbacks may work in both direction however. Recent geomorphological and thermochronometric work indicates that the location of active faulting in mountains and thus orogenic architecture is controlled by climate zonation (Hodges et al. 2004; Wobus et al. 2003), at least as much as by plate tectonic processes. Both wedge and channel-flow models proposed to explain the origin of the Greater Himalaya (Nelson et al. 1996; Robinson et al. 2003; Harris 2007) require focused erosion driven by monsoon rains to allow exhumation of deeply buried metamorphic rocks. If we are to understand what influence climate alone has on erosion (and thus on tectonics) then we need to isolate the climate signal from tectonic overprints. One way to do this is to study changes in erosion on timescales that are too short to allow tectonism to be an important influence. In this study we assess the response of the Indus river basin to the intensification of the South Asian monsoon since the Last Glacial Maximum (LGM), around 20 ka. Cave records in Arabia (Fleitmann et al. 2003), together with Indian lake sediments (Enzel et al. 1999) and marine cores (Staubwasser et al. 2002) now show that the summer monsoon strengthened as global climate warmed between the LGM (c. 20 ka) and the Early Holocene c. 8 ka. Earlier sedimentological work in the Bengal delta demonstrated that the intensification of the summer monsoon correlated with a great increase in the rate of sediment delivery to the delta during the Early Holocene, c. 8 ka (Goodbred & Kuehl 2000). This is suggestive of greatly enhanced erosion and sediment transport rates in the source regions, presumably driven by the influence of stronger summer monsoon rains, as the speed of response rules out a tectonic trigger. Indeed studies of landsliding in the western Himalaya show that large scale mass wasting correlates with periods of stronger monsoon (Bookhagen et al. 2005), consistent with the delta records. Recently we reported that the Indus delta also prograded seawards during the onset of the Holocene (Giosan et al. 2006), that is, at a time of rapid sea level rise. This observation requires a major increase in the flux of sediment to the delta to balance the increasing accommodation space caused by sea level rise. Crucially, the delta was prograding southwards during the 8–12 ka period when rates of eustatic sea level rise were greatest (Camoin et al. 2004). Most recently, Clift et al. (2008) used a combination of bulk sediment Nd

isotopes, single grain U –Pb zircon ages and Ar –Ar muscovite mica ages to argue for a sharp increase in the relative erosional flux from the Lesser Himalaya during the Early Holocene. In this paper we test their hypothesis that intensification of the summer monsoon resulted in much heavier rain along the southern edge of the Greater Himalaya and especially over the Lesser Himalaya. We do this using a series of additional geochemical proxies to assess the source of sediment at any given time and the changing environmental conditions in the Indus basin, which can then be compared to established climate records.

Sampling Samples were taken from four boreholes drilled in the delta (Fig. 1). Although total recovery was not high material was recovered from most parts of the drilled sections, thus allowing a relatively continuous erosion record to be reconstructed (Fig. 2). At Keti Bandar drilling penetrated the ravinement surface that forms the base of the modern Indus delta and recovered Pleistocene sand deposited during the LGM (Clift et al. 2008). At all other sites only the Holocene to Younger Dryas sections were recovered. Ages of deposition were calculated from accelerator mass spectrometer (AMS) 14C dating of organic materials made at the National Ocean Sciences Accelerator Mass Spectrometry facility (NOSAMS, Woods Hole Oceanographic Institution) (Clift et al. 2008). Radiocarbon ages from below the deltaic sediments were 28.7 and 38.9 ka, suggesting reworking and mixing of older sediment prior to transgression after 20 ka. The Keti Bandar section shows two coarsening-upward cycles, separated by a transgressive mud deposited after c. 8 ka (Fig. 2). The cores were sampled both for sands, which were mostly used for single grain thermochronological methods, and for clays that were used for organic carbon analysis.

Analytical methods A number of different provenance methods were used in order to establish a matrix of constraints, since typically one or two methods were insufficient to define a sediment source area for any given sand sample. In this study, we analysed the sediments using classical sand petrography, Sr isotopes, apatite and zircon fission track, Ar –Ar single biotite grain dating and additional U –Pb zircon dating, beyond the samples already considered by Clift et al. (2008). In addition, we selected clay samples for organic carbon isotope geochemistry in order to constrain the general composition of the vegetation in the Indus basin. Here we describe

MONSOON CONTROL OVER EROSION

HK

187

KK K NP

Tarbela Dam

LA 35° N

ZA

Islamabad m Ch

llu

vi Ra

ks

Je

GH ali Siw

b

a en

STD

LH MBT

MCT Garhwal MFT

Sutlej

30° N

us

Ind

rt

Sukkur

ar

De

se

Thatta

Th

Karachi

Keti Bandar

Gularchy Jati 25° N

Arabian Sea 70°E

75°E

Fig. 1. Shaded topographic map of the Indus drainage basin, showing the location of the boreholes in the delta, the major tributaries of the modern river and the principal sediment source terrains (HK, Hindu Kush; K, Kohistan; NP, Nanga Parbat; ZA, Zanskar; LA, Ladakh; KK, Karakoram; GH, Greater Himalaya; LH, Lesser Himalaya; MBT, Main Boundary Thrust; MFT, Main Frontal Thrust; MCT, Main Central Thrust; STD, South Tibet Detachment) [from Clift et al. (2008)]. White dots indicate boreholes. Black dots with white rim indicate place name mentioned in the text. Reproduced with the permission of the Geological Society of America.

the methods employed to reconstruct the evolving character of the Indus delta.

Mineralogy All nine samples analysed for petrography from the Keti Bandar and Thatta cores were very finegrained sands containing variable amounts of bioclasts. In each sample, 400 points were counted by the Gazzi-Dickinson method (Ingersoll et al. 1984). Data are presented in Tables 1 and 2. Traditional ternary parameters and plots (Dickinson 1985) were supplemented, specifically as lithic

grains are concerned, by an extended spectrum of key indices. Metamorphic rock fragments were classified according to both composition and metamorphic rank, mainly inferred from degree of recrystallization of mica flakes (Garzanti & Vezzoli 2003). Very low to low-rank metamorphic lithics, for which protolith can still be inferred, were subdivided into metasedimentary (Lms) and metavolcanic (Lmv) categories. Medium to high-rank metamorphic lithics were subdivided instead into felsic (metapelite, metapsammite, metafelsite; Lmf) and mafic (metabasite; Lmb) categories.

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Fig. 2. Diagram shows the stratigraphy of the sediment cored at each of the drilling sites considered in this study, as shown in Figure 1. Each log shows the depth, proportion of the section recovered by drilling (black equates to recovery), the age control points constrained by 14C AMS dating of organic materials, as well as the location of samples selected from isotopic and thermochronological analysis. TD, total depth. Reproduced with the permission of the Geological Society of America.

Heavy minerals were separated from the veryfine to fine sand fraction (63–250 mm), treated with acetic acid and sodium dithionite, by centrifuging in sodium metatungstate (density c. 2.9 g cm23). From each sample, 200– 250 transparent heavy minerals were counted in grain mounts by the ‘ribbon-counting’ or ‘Fleet’ methods (Mange & Maurer 1992). The abundance of total (transparent þ opaque þ turbid) and transparent heavy minerals in the sediment were expressed by the ‘Heavy Mineral Concentration’ (HMC) and ‘transparent Heavy Mineral Concentration’ (tHMC) indices (Garzanti & Ando` 2007).

Sr isotope analysis Samples from both the Keti Bandar and Gularchy boreholes were sampled for Sr analysis. Sr analysis was only performed on samples that had previously been analysed for Nd isotopes (Clift et al. 2008). Sr can be used for provenance work, but because

it is mobilized during chemical weathering is less definitive than Nd isotopes. However, this characteristic can be used to effect because Sr variability in the absence of provenance changes has been interpreted to reflect changing intensities of chemical weathering (Palmer & Edmond 1991; Blum & Erel 1995). In particular, the Sr isotope composition of clay reflects the Sr characteristics of the groundwater in which they were formed, rather than the composition of the source rock (McBride 1994; Derry & France-Lanord 1996). Assuming that the clays have not been affected by diagenesis (unlikely in such young materials) the Sr isotopes can be used to track the evolving intensity of silicate weathering. However, in this setting the Nd isotope composition is known to change especially at 14 –8 ka and so changes in Sr at those times may be mostly provenance driven. Samples were accurately weighed into teflon screw-top beakers and dissolved using HF-HNO3HCl. Sr samples were separated in 2.5 N HCl

Table 1. Bulk mineralogy data for selected sands from the Keti Bandar and Thatta boreholes derived from point counting of microcopic grain mounts Site

n/a KB-5-2 TH-10-8 KB-20-1 KB-23-2 KB-26-2 KB-30-1 KB-34-4 KB-40-5 KB-41-2

Delta Keti Bandar Thatta Keti Bandar Keti Bandar Keti Bandar Keti Bandar Keti Bandar Keti Bandar Keti Bandar

Depth (m)

Age (ka)

Mean grain size (micron)

Q

KF

P

15 31 58 67 76 88 102 118 120

0 0.18 7.14 9.21 9.65 10.64 12.26 14.23 20.00 20.00

119 82 93 87 39 51 38 73 85 82

41 40 38 46 40 36 37 41 39 46

7 10 11 10 10 8 11 5 12 8

9 12 15 10 10 8 8 15 14 14

Lvf Lvm Lcc Lcd Lp Lch Lms Lmv Lmf Lmb Lu Mu Bi

0 0 1 1 0 0 2 1 1 1

0 0 0 0 0 0 0 0 1 1

8 8 6 11 6 7 7 6 0 0

3 3 3 3 1 2 2 2 4 4

3 1 1 1 2 0 1 5 1 0

1 0 0 0 0 0 0 1 0 0

7 3 4 4 4 2 2 5 4 6

1 1 4 1 1 1 1 4 2 2

3 5 5 5 5 2 3 4 5 5

2 0 1 0 0 0 0 0 1 0

0 0 0 0 0 0 0 0 0 0

1 4 2 2 4 7 6 4 4 2

2 11 8 5 15 25 19 8 9 7

Dense Total minerals 11 3 2 2 1 2 1 0 3 2

100 100 100 100 100 100 100 100 100 100

MONSOON CONTROL OVER EROSION

Sample

Abbreviations: Q, quartz; KF, K-felspar; P, plagioclase; Lvf, lithic felsic volcanic; Lvm, lithic mafic volcanic; Lcc, calcite carbonate lithic; Lcd, dolomite carbonate lithic; Lp, shale/silt; Lch, chert; Lmv, metavolcanic lithics; Lmf, felsic metamorphic lithics; Lmb, basic metamorphic lithics; Lu, ultramafic lithics; Mu, muscovites; Bi, biotites.

189

190

Table 2. Heavy mineral data for selected sands from the Keti Bandar and Thatta boreholes derived from point counting of microcopic grain mounts Sample

Depth (m)

Age (ka)

% of heavy minerals in 63 – 250 mm fraction

% of heavy minerals transparent

Delta Keti Bandar Thatta Keti Bandar Keti Bandar Keti Bandar Keti Bandar Keti Bandar Keti Bandar Keti Bandar

n/a KB-5-2 TH-10-8 KB-20-1 KB-23-2 KB-26-2 KB-30-1 KB-34-4 KB-40-5 KB-41-2

0 14 30 58 67 77 78 101 118 120

0 0.18 7.14 9.21 9.65 10.64 10.77 14.20 20.00 20.00

9.6 3.0 4.9 3.9 0.9 1.3 1.4 4.0 4.3 4.3

8.0 2.7 3.9 3.4 0.7 1.1 1.2 3.2 3.8 3.6

% % % Total zircon tourmaline rutile sphene transparent opaque turbid

83 91 79 87 83 88 84 81 88 84

10 1 6 3 2 1 4 3 5 3

8 9 15 10 15 11 12 16 8 13

100 100 100 100 100 100 100 100 100 100

0 0 0 1 0 0 0 0 0 2

2 3 2 3 2 2 2 3 2 0

1 1 0 1 2 0 1 0 0 0

bluegreen greenbrown/ green hornblende brown red hornblende hornblende hornblende

2 2 0 3 2 4 1 1 3 1

39 45 33 39 34 38 34 36 47 51

4 1 5 2 4 5 4 2 0 4

3 2 2 4 0 1 2 1 3 2

2 0 0 0 0 0 0 0 0 0

Site

Sample

tremolite

actinolite

green augite

diopside/ hedembergite

enstatite

hypersthene

epidote

clinozoisite

zoisite

other epidotes

chloritoid

garnet

staurolite

kyanite

sillimanite

Total

Delta Keti Bandar Thatta Keti Bandar Keti Bandar Keti Bandar Keti Bandar Keti Bandar Keti Bandar Keti Bandar

n/a KB-5-2 TH-10-8 KB-20-1 KB-23-2 KB-26-2 KB-30-1 KB-34-4 KB-40-5 KB-41-2

0 0 0 0 2 2 1 0 0 0

2 3 8 6 13 7 8 5 3 4

0 0 0 0 0 0 0 2 0 0

2 4 7 5 4 4 8 4 0 1

0 0 0 0 0 0 1 0 0 0

1 3 2 2 1 1 1 1 0 1

21 28 22 26 19 25 28 25 32 24

1 1 1 1 1 0 0 0 0 0

3 0 0 1 0 2 0 1 1 1

0 0 0 1 0 0 0 1 1 0

0 0 0 2 1 0 0 0 0 0

12 2 13 4 5 5 3 11 4 4

1 1 0 0 1 1 0 1 0 0

1 1 1 3 3 1 0 0 0 1

1 0 1 1 2 1 0 2 0 0

100 100 100 100 100 100 100 100 100 100

P. D. CLIFT ET AL.

Site

MONSOON CONTROL OVER EROSION

using Bio-Rad AG50W X8 200-400 mesh cation exchange resin using standard column methods. They were analysed on a by Finnigan ‘Neptune’ multi-collector inductively coupled plasma mass spectrometer (MC-ICP-MS) at Woods Hole Oceanographic Institution. Sample measurements were normalized to 86Sr/88Sr ¼ 0.1194 and referenced to a value of 0.710240 for NBS987 standard. Results are provided in Table 3.

Organic carbon analysis Total organic carbon content (TOC) and stable isotopic composition of the bulk organic matter (d13C) were determined for 20 samples taken exclusively from the Keti Bandar borehole (Fig. 1). A detailed description of the analytical method can be found in Galy et al. (2007a). Indus delta sediments contain significant amounts of detrital carbonates including dolomite. TOC and isotopic measurements must therefore be performed on decarbonated sediments. Efficient dolomite dissolution was achieved through 1 hr leaching with 4 wt% HCl at 80 8C. After 50 8C oven drying, the organic carbon content and stable isotopic composition were determined by EA-IRMS (elemental analysis – isotope ratio mass spectrometry). organic carbon solubilization during acid treatment was taken into account following the approach described by Galy et al. (2007a). The overall 2s uncertainty associated with the TOC and d13C determination is respectively 0.02% and 0.25‰. Results of the carbon isotope analysis are provided in Table 4.

Major and trace element analysis Major and trace element concentrations for the muds which were also measures for carbon isotopes were measured respectively by Inductively Coupled Plasma Atomic Emission Spectrophotometry (ICP-AES) and Inductively Coupled Plasma Mass Spectrometry (ICP-MS) at the Service d’Analyse des Roches et des Mine´raux (CRPG-Nancy, France) on bulk sediment after lithium metaborate fusion. Results of the element analysis are provided in Table 4.

Fission track analyses In this study we employed the fission-track method applied to apatite, which records cooling through c. 125 –60 8C over timescales of 1– 10 Ma and to zircon which records cooling through c. 200 8C (Green 1989). Fission-track analysis was performed at University College, London, UK using five samples, one of which, dating from the LGM, was analysed for both apatite and zircon. Polished grain mounts were etched with 5N HNO3 at 20 8C

191

for 20 seconds to reveal the spontaneous fission tracks. Subsequently, the uranium content of each crystal was determined by irradiation, which induced fission of 235U. The induced tracks were registered in external mica detectors. The samples for this study were irradiated in the thermal facility of the Hifar Reactor at Lucas Heights, Australia. The neutron flux was monitored by including Corning glass dosimeter CN-5, with a known uranium content of 11 ppm, at either end of the sample stack. After irradiation, sample and dosimeter mica detectors were etched in 48% HF at 20 8C for 45 minutes. Only crystals with sections parallel to the c-axis were counted, as these crystals have the lowest bulk etch rate. To avoid biasing results through preferred selection of apatite crystals, the samples were systematically scanned and each crystal encountered with the correct orientation was analysed, irrespective of track density. The results of the fission-track analysis are presented in Table 5.

Ar –Ar mica dating Single crystal 40Ar/39Ar laser-fusion analyses were performed on biotite grains separated from two sands (KB-41-2, from the LGM deposits at Keti Bandar, and TH-4-6 from c. 6.4 ka sand at Thatta) at the Massachusetts Institute of Technology (MIT). Biotites from a modern sample of the Indus at Thatta were previously published by Clift et al. (2004). Prior to analysis, samples were irradiated in the C5 position of the McMaster University Nuclear Reactor, Canada, using 1 mm Cd shielding for four hours at a power level of 2 MW. After fusion with an Ar-ion laser, the released gases were purified for 10 minutes with two Al –Zr getters operated at 400 8C and room temperature, respectively, and then admitted to an MAP 215-50 mass spectrometer for Ar isotopic analysis using a Johnson MM-1 electronic multiplier operated at a gain of about 10 000. The conversion efficiency of 39K to 39Ar was monitored using sanidine from the Taylor Creek rhyolite (TCR-2a) assuming an age of 28.34 Ma (Renne et al. 1998), and is known to better than 0.3% (1s). Corrections for neutron-induced interferences, determined using Fe-doped kalsilite glass and optical CaF2, were 0.00039 for 40Ar/39ArK, 0.01243 for 38Ar/39ArK, 0.000672 for 39Ar/37ArCa, 0.000033 for 38Ar/ 37 ArCa and 0.00028 for 36Ar/37ArCa. Final data reduction was conducted with the program ArArCalc (Koppers 2002); results are shown in Table 6.

U – Pb zircon dating U –Th –Pb isotopic compositions of zircon grains were analysed at the Australian National University,

192

Table 3. Sr isotopic data from Keti Bandar and Gularchy boreholes. Matching Nd isotope data and 14C ages are from Clift et al. (2008). See Figure 1 for locations Sample number

Fine Sand Fine Sand Clay Clay Silt Clay Fine Sand Sand Sand Sand

Location

Depth (m)

Keti Bandar Keti Bandar Keti Bandar Keti Bandar Keti Bandar Keti Bandar Keti Bandar Gularchy Gularchy Gularchy

14.0 55.0 61.0 92.0 97.0 108.0 117.0 11.3 26.2 47.5

14

C Age (ka)

0.21 8.72 9.42 12.20 12.91 13.73 28.70 2.65 6.42 10.94

Age 1s (ka)

Nd143/Nd144

1Nd

0.00 0.15 0.20 0.32 0.32 0.37 0.00 0.08 0.08 0.18

0.511947 0.511858 0.511986 0.511919 0.512033 0.512000 0.512082 0.511911 0.511931 0.511962

213.5 215.2 212.7 214.0 211.8 212.4 210.8 214.2 213.8 213.2

87

Sr/86Sr

0.724512 0.727689 0.719290 0.719951 0.710671 0.714531 0.714764 0.721362 0.719811 0.725178

P. D. CLIFT ET AL.

KB-5-2 KB-19-4 KB-21-1 KB-31-2 KB-34-4 KB-37-4 KB-40-5 GUL-ZP-5-1 GUL-ZP-10-2 GUL-ZP-17-1

Lithology

Table 4. Carbon isotope and organic carbon data, together with associated major and trace element compositions from Holocene samples from Keti Bandar and Gularchy boreholes. See Figure 1 for locations Sample # Depth Age SiO2 Al2O3 Fe2O3 MnO MgO CaO Na2O K2O TiO2 P2O5 (m) (a bp) (%) (%) (%) (%) (%) (%) (%) (%) (%) (%) 7.5 20.0 26.0 26.5 37.0 44.0 55.0 59.0 67.0 70.5 92.0 96.0 100.0 100.5 107.0 110.0 112.0 117.0 118.0 119.0

112 299 1235 1314 4086 6336 8140 8560 9400 9768 12026 12446 12866 12918 13601 13916 14126 14651 14756 28700

54.23 55.93 51.48 51.16 54.02 50.67 53.44 65.28 54.32 63.43 54.09 51.49 50.27 50.67 52.36 49.58 51.51 50.89 48.04 50.07

13.47 13.42 15.04 15.22 13.69 14.84 13.67 10.94 13.72 11.04 13.60 13.78 14.13 14.49 14.04 14.84 14.18 14.58 14.86 15.23

5.46 5.35 6.42 6.51 5.64 6.38 5.62 4.00 5.67 4.79 5.66 6.11 6.04 6.21 5.88 6.42 6.11 6.35 6.68 6.63

0.09 0.09 0.10 0.10 0.09 0.10 0.10 0.07 0.10 0.08 0.09 0.09 0.09 0.09 0.09 0.09 0.09 0.10 0.10 0.10

2.94 2.87 3.32 3.39 3.04 3.42 3.00 2.22 3.04 2.35 3.03 3.20 3.46 3.38 3.48 3.45 3.56 3.61 4.36 3.49

8.48 7.27 7.28 7.17 7.92 7.71 8.14 6.72 7.26 6.81 8.38 8.76 9.39 8.28 8.30 8.64 8.09 7.91 7.67 7.69

1.74 2.14 1.83 1.88 1.83 1.73 1.79 1.87 2.35 1.94 1.82 1.90 1.91 1.87 2.02 1.81 1.96 1.98 1.88 1.74

2.54 2.54 2.94 2.98 2.53 2.88 2.64 2.09 2.49 2.08 2.62 2.68 2.71 2.83 2.76 2.89 2.82 2.89 2.95 2.97

0.66 0.69 0.74 0.75 0.81 0.72 0.71 0.60 0.71 0.63 0.68 0.69 0.70 0.71 0.70 0.71 0.70 0.70 0.72 0.72

0.14 0.16 0.16 0.16 0.19 0.15 0.15 0.14 0.15 0.16 0.15 0.16 0.16 0.15 0.16 0.15 0.16 0.16 0.14 0.15

Total (%)

10.05 9.39 10.38 10.68 10.11 11.04 10.29 6.56 10.19 7.08 10.29 11.12 11.67 11.34 10.68 11.70 10.43 11.75 12.80 11.41

99.79 99.86 99.67 99.98 99.86 99.64 99.55 100.48 100.00 100.39 100.39 99.99 100.53 100.00 100.46 100.28 99.61 100.92 100.20 100.21

As Ba Be Bi Cd Ce Co (ppm) (ppm) (ppm) (ppm) (ppm) (ppm) (ppm) 9.8 10.9 13.1 12.3 8.6 9.9 10.2 5.2 14.1 3.9 12.3 19.4 12.0 14.6 13.0 17.1 13.5 12.2 8.4 11.9

348.0 363.9 379.2 384.7 329.5 370.2 370.2 328.4 425.9 325.9 361.0 364.6 364.8 367.9 375.3 378.6 371.9 364.1 337.3 380.7

2.4 2.5 2.8 2.7 2.4 2.7 2.7 2.4 2.5 2.2 2.4 2.3 2.5 2.4 2.3 2.6 2.4 2.4 2.4 2.6

0.9 0.8 1.0 1.0 0.8 1.0 1.0 0.6 0.5 0.6 0.5 0.6 0.6 0.7 0.6 0.7 0.6 0.6 0.5 0.7

,L.D. ,L.D. ,L.D. ,L.D. 0.287 ,L.D. ,L.D. ,L.D. ,L.D. 0.271 ,L.D. ,L.D. ,L.D. ,L.D. ,L.D. ,L.D. ,L.D. ,L.D. ,L.D. ,L.D.

60.8 71.7 68.0 65.8 92.7 64.5 72.3 66.2 66.2 75.0 62.5 62.1 62.6 62.2 61.3 66.2 61.4 61.6 62.7 65.3

14.8 13.9 16.2 16.7 14.8 16.9 15.5 10.9 15.5 10.5 15.2 15.5 16.2 16.1 15.6 17.8 17.1 17.6 16.9 17.2

MONSOON CONTROL OVER EROSION

KB-3-3 KB-8-2 KB-10-2 KB-10-5 KB-13-2 KB-16-1 KB-19-3 KB-21-2 KB-23-4 KB-25-2 KB-31-1 KB-33-3 KB-34-2 KB-35-2 KB-37-4 KB-38-2 KB-38-3 KB-40-2 KB-40-3 KB-41-1

PF (%)

(Continued )

193

194

Table 4. Continued Cr Cs Cu Dy Er Eu Ga Gd Ge Hf Ho In La Lu Mo Nb Nd Ni Pb Pr Rb (ppm) (ppm) (ppm) (ppm) (ppm) (ppm) (ppm) (ppm) (ppm) (ppm) (ppm) (ppm) (ppm) (ppm) (ppm) (ppm) (ppm) (ppm) (ppm) (ppm) (ppm)

KB-3-3 KB-8-2 KB-10-2 KB-10-5 KB-13-2 KB-16-1 KB-19-3 KB-21-2 KB-23-4 KB-25-2 KB-31-1 KB-33-3 KB-34-2 KB-35-2 KB-37-4 KB-38-2 KB-38-3 KB-40-2 KB-40-3 KB-41-1

100.9 92.2 107.6 107.9 112.9 110.4 105.2 83.1 97.8 84.2 99.0 104.4 105.2 108.7 108.3 114.7 111.6 113.0 121.8 114.7

8.2 7.8 10.1 10.0 8.2 10.1 9.3 6.2 6.9 5.9 8.3 8.4 7.8 8.9 7.8 9.1 8.2 8.7 8.2 10.7

25.6 24.4 31.4 33.5 24.8 34.0 27.8 13.8 27.1 15.5 26.4 28.5 31.0 32.3 28.7 37.4 30.5 33.6 52.6 33.6

4.0 4.5 4.3 4.1 5.9 4.1 4.8 4.3 4.4 4.7 4.0 4.0 4.0 4.0 4.0 4.1 4.1 4.0 4.2 4.1

2.2 2.5 2.4 2.3 3.2 2.2 2.6 2.4 2.4 2.5 2.1 2.2 2.2 2.2 2.2 2.2 2.2 2.2 2.3 2.3

1.1 1.2 1.1 1.1 1.4 1.1 1.2 1.1 1.1 1.2 1.0 1.1 1.1 1.1 1.1 1.1 1.1 1.1 1.1 1.1

17.5 17.5 20.3 20.0 18.7 20.2 19.3 14.4 18.5 14.4 17.8 18.4 18.7 19.1 17.5 19.5 18.6 19.3 20.0 20.4

4.4 5.0 4.7 4.6 6.5 4.5 5.2 4.8 4.9 5.3 4.5 4.5 4.4 4.4 4.5 4.7 4.6 4.5 4.6 4.6

1.6 1.6 1.7 1.7 1.7 1.7 1.7 1.5 1.6 1.6 1.6 1.6 1.6 1.5 1.6 1.6 1.6 1.6 1.7 1.7

3.6 5.1 4.0 3.8 7.7 3.5 4.7 6.1 4.5 6.8 4.1 4.1 3.9 3.8 4.0 3.6 4.2 4.0 3.6 3.5

0.8 0.9 0.8 0.8 1.1 0.8 0.9 0.8 0.8 0.9 0.8 0.8 0.8 0.8 0.8 0.8 0.8 0.8 0.8 0.8

0.2 0.2 0.2 0.2 0.3 0.2 0.3 0.2 0.2 0.2 0.2 0.2 0.3 0.2 0.1 0.2 0.2 0.2 0.2 0.2

30.2 36.3 33.9 33.4 46.5 31.9 36.1 33.0 33.6 37.8 31.6 31.3 31.6 31.5 31.1 33.1 31.0 31.2 31.5 32.8

0.3 0.4 0.4 0.4 0.5 0.3 0.4 0.4 0.4 0.4 0.3 0.3 0.3 0.3 0.3 0.4 0.3 0.3 0.4 0.4

0.6 ,L.D. 0.7 0.7 0.5 0.8 0.6 ,L.D. 0.7 ,L.D. 0.6 1.3 0.7 0.9 0.6 0.7 0.6 0.6 0.8 0.7

12.2 13.0 14.2 14.4 15.1 13.4 13.8 11.3 13.9 11.6 12.7 12.9 13.1 12.9 12.5 13.4 12.8 13.0 13.2 13.8

26.2 30.9 28.6 28.5 40.1 27.5 31.0 28.6 28.6 32.3 26.5 26.8 26.7 26.8 26.5 28.3 26.4 26.9 27.3 27.7

53.5 48.1 60.7 59.2 53.3 61.0 54.4 36.9 53.9 34.4 54.7 55.6 59.4 62.3 56.6 64.9 61.4 65.8 71.8 62.9

17.9 16.7 22.1 20.6 19.1 20.3 18.4 19.9 11.1 17.6 18.1 15.9 14.7 16.6 14.4 15.1 15.8 14.9 14.0 20.5

6.9 8.3 7.7 7.6 10.6 7.3 8.3 7.7 7.7 8.7 7.2 7.2 7.2 7.1 7.1 7.6 7.1 7.1 7.2 7.5

117.7 117.0 141.6 136.1 119.7 133.7 127.8 97.8 115.7 95.0 119.8 121.1 121.2 125.8 120.1 131.7 123.2 125.8 124.2 143.0

(Continued )

P. D. CLIFT ET AL.

Sample #

Table 4. Continued Sb (ppm)

Sm (ppm)

Sn (ppm)

Sr (ppm)

Ta (ppm)

Tb (ppm)

Th (ppm)

Tm (ppm)

U (ppm)

V (ppm)

W (ppm)

Y (ppm)

Yb (ppm)

Zn (ppm)

Zr (ppm)

TOC (%)

d13C (‰)

KB-3-3 KB-8-2 KB-10-2 KB-10-5 KB-13-2 KB-16-1 KB-19-3 KB-21-2 KB-23-4 KB-25-2 KB-31-1 KB-33-3 KB-34-2 KB-35-2 KB-37-4 KB-38-2 KB-38-3 KB-40-2 KB-40-3 KB-41-1

0.7 0.7 0.7 0.7 0.7 0.7 0.7 0.5 1.7 0.5 1.6 1.7 1.9 0.8 0.7 2.0 1.7 1.7 1.6 0.7

5.2 6.0 5.6 5.5 7.9 5.4 6.2 5.8 5.7 6.4 5.3 5.2 5.4 5.2 5.3 5.6 5.3 5.3 5.3 5.5

3.6 3.8 4.2 4.2 3.8 4.2 4.0 3.4 4.4 3.4 3.9 4.0 4.2 3.7 3.5 4.3 4.1 4.0 3.9 4.0

204.6 201.1 203.2 194.4 214.3 190.9 199.0 188.8 199.7 191.3 210.2 253.2 370.8 221.1 223.1 226.0 233.6 230.2 181.4 198.7

1.1 1.3 1.3 1.3 1.5 1.2 1.2 1.2 1.3 1.2 1.1 1.2 1.2 1.2 1.2 1.2 1.2 1.2 1.2 1.3

0.7 0.8 0.7 0.7 1.0 0.7 0.8 0.8 0.8 0.8 0.7 0.7 0.7 0.7 0.7 0.7 0.7 0.7 0.7 0.7

12.8 15.3 15.5 15.4 20.9 14.3 15.7 13.4 14.9 15.6 13.6 14.0 14.0 14.1 13.8 15.1 13.7 14.0 14.2 15.2

0.3 0.4 0.4 0.3 0.5 0.3 0.4 0.4 0.4 0.4 0.3 0.3 0.3 0.3 0.3 0.3 0.3 0.3 0.3 0.3

2.3 2.9 2.7 2.6 3.8 2.5 2.9 2.5 2.9 2.9 2.5 2.6 2.6 2.7 2.6 2.7 2.7 2.7 2.8 2.6

100.9 96.0 115.6 115.2 104.1 115.7 105.5 71.6 102.3 70.8 101.1 106.0 110.5 115.9 106.6 119.9 110.6 114.1 125.5 119.0

2.3 2.9 2.8 2.8 2.7 2.6 2.6 3.5 2.7 10.5 2.3 2.4 2.4 2.4 2.3 2.5 2.3 2.4 2.4 2.7

23.4 25.6 24.3 23.7 33.4 23.0 27.4 25.3 25.0 26.6 23.1 22.9 22.4 22.0 23.2 23.6 23.2 22.6 23.5 23.2

2.2 2.5 2.4 2.3 3.3 2.2 2.6 2.4 2.4 2.6 2.2 2.2 2.2 2.2 2.3 2.3 2.3 2.3 2.4 2.3

82.7 87.3 120.1 94.9 94.2 100.1 95.5 63.9 122.4 62.8 82.1 88.1 89.8 88.8 85.5 94.5 89.2 91.0 95.5 98.1

133.2 187.9 143.9 138.4 291.0 126.9 170.9 224.7 164.6 268.0 152.5 152.5 143.1 131.7 143.6 126.3 150.7 144.2 130.0 125.3

0.27 0.26 0.37 0.34 0.32 0.43 0.31 0.16 0.31 0.18 0.31 0.31 0.32 0.29 0.26 0.32 0.29 0.26 0.26 0.44

221.2 220.5 219.8 220.3 219.8 220.0 220.3 222.6 221.0 223.3 220.7 221.1 220.9 221.7 220.8 221.5 220.5 220.8 220.0 220.0

MONSOON CONTROL OVER EROSION

Sample #

195

196

Table 5. Fission track analytical data for Holocene Indus delta sands Sample no./ Strat age/ No. of Field no. lithology crystals

No apatite Apatite Apatite Apatite Apatite Zircon Apatite

44 15 16 70 28 61

rd

Nd

1.317 1.317 1.220 1.317 0.493 1.220

3651 3651 6107 3651 3450 6107

Spontaneous rs

Ns

Induced ri

Ni

0.101 135 5.538 6762 0.221 57 3.681 949 0.109 50 4.036 1857 0.162 661 4.122 16735 6.107 2404 5.414 2131 0.846 297 4.05 14219

Age dispersion Px2 RE% 0 3.0 5.3 0 0 0

Central age (Ma) + 1s

1st age comp.

2nd age comp.

3rd age comp.

4th age comp.

124.5 4.4 + 1.0 2.6 + 0.3 (40) 9 + 3 (1) 24 + 4 (3) 56.7 11.1 + 2.7 7.4 + 1.7(13) 23 + 4 (2) 40.4 5.5 + 1.0 1.8 + 0.7 (6) 7.7 + 1.2 (10) 45.5 8.9 + 0.7 3.5 + 0.4 (20) 8.7 + 0.5 (50) 64.7 34.9 + 4.4 17.1 + 0.9 (14) 42.3 + 4.2 (9) 96 + 11 (4) 79.8 4.4 + 0.6 2.2 + 0.3 (40) 9.6 + 0.9 (21)

Note: (i) Track densities are (106 tr cm22) numbers of tracks counted (N) shown in brackets. Rd is the induced track density of the dosimeter. Nd is the number of tracks counted in the dosimeter/ Ns is the number of spontaneous tracks counted in the sample. Ni is the number of induced tracks counted in the sample. (ii) Analyses by external detector method using 0.5 for the 4p/2p geometry correction factor; (iii) Ages calculated using dosimeter glass CN-5; (apatite) zCN5 ¼338 + 4; CN-2 (zircon) zCN2 ¼127+4 calibrated by multiple analyses of IUGS apatite and zircon age standards (Hurford 1990); (iv) Px2 is the probability of obtaining x2 value for v degrees of freedom, where v ¼ no. crystals – 1. RE %, age dispersion (or the spread of the individual crystal data) is given by the % relative standard deviation of the central age; (v) Central age is a modal age, weighted for different precisions of individual crystals (Galbraith 1990).

P. D. CLIFT ET AL.

KB-34-4 KB-19-4 KB-23-3 KB-5-2 KB40-5 KB40-5 TH-10-8

Dosimeter

MONSOON CONTROL OVER EROSION

197

Table 6. Ar –Ar analytical data for detrital biotite grains extracted from Holocene Indus delta sands KB-41-2 Age (Ma) 6.88 7.45 7.70 8.05 8.69 9.09 9.09 9.14 9.32 9.35 9.61 9.67 9.80 10.16 10.26 10.49 11.37 11.52 12.69 15.40 15.87 15.89 16.30 16.44 18.11 18.51 19.55 21.90 24.96 25.52 26.45 26.52 26.55 30.15 32.80 33.23 38.04 40.60 42.28 43.18 43.80 116.42 126.38 246.18 366.70 572.73 TH-4-6 Age (Ma) 0.22 0.41 0.48 0.82 0.97 1.29 1.36 1.39 1.71

+2s +1.64 +3.23 +0.44 +2.80 +0.66 +0.37 +0.21 +0.44 +0.18 +0.81 +3.81 +0.83 +0.39 +0.93 +3.69 +0.70 +0.65 +2.31 +1.71 +2.41 +0.79 +1.04 +0.87 +0.64 +0.51 +0.34 +0.74 +3.90 +1.98 +1.02 +1.48 +1.07 +0.54 +0.67 +0.55 +0.78 +1.14 +0.60 +1.07 +0.57 +0.39 +0.81 +0.93 +1.16 +1.99 +3.19 +2s +1.07 +0.91 +0.55 +0.67 +1.13 +0.37 +0.46 +0.64 +0.75

40

Ar(r) (%)

12.14 6.28 35.45 4.70 55.30 55.48 77.40 48.89 84.63 84.66 5.08 59.20 69.54 66.19 6.06 82.14 77.99 16.00 13.29 10.24 46.97 41.46 25.53 76.76 58.42 82.14 86.39 8.48 20.80 81.10 32.09 89.69 86.09 61.70 92.82 90.30 42.20 85.08 53.26 58.88 78.45 95.22 89.48 98.97 98.92 99.29 40

Ar(r) (%)

8.34 9.58 5.14 13.40 13.22 46.53 22.69 28.52 20.85

39

Ar(k) (%)

TH-4-6 Age (Ma)

1.35 0.74 3.42 3.35 1.84 3.91 2.79 3.88 4.17 1.30 1.00 1.35 2.82 1.20 0.75 1.43 1.63 0.83 1.67 2.34 1.83 1.51 2.88 1.59 3.26 2.25 1.40 2.13 2.06 1.18 1.65 1.04 2.12 2.37 1.97 1.67 2.07 1.83 1.88 4.75 5.36 1.92 1.76 3.60 1.47 2.68

1.85 2.17 2.37 2.53 2.56 2.75 2.76 2.76 2.76 2.79 3.22 3.24 3.24 3.31 3.32 3.32 3.38 3.42 3.44 3.53 3.54 3.55 3.66 3.75 3.84 3.87 3.97 4.17 4.34 4.42 4.49 4.5 4.51 4.53 4.58 4.63 4.63 4.71 4.76 4.83 4.86 4.91 5.03 5.04 5.10 5.14 5.15 5.18 5.20 5.34 5.57 5.87 6.24 6.27 6.30 6.35 6.52 6.55 6.70

39

Ar(k) (%)

0.25 0.29 0.56 0.43 0.22 0.66 0.87 0.43 0.37

+2s +0.60 +0.94 +0.33 +0.54 +0.36 +0.56 +0.24 +0.63 +0.23 +0.75 +0.30 +0.79 +0.88 +0.64 +0.56 +0.52 +0.20 +0.84 +0.29 +0.69 +0.27 +0.59 +0.44 +0.56 +0.45 +0.37 +0.47 +0.45 +0.25 +0.67 +0.39 +0.50 +0.36 +0.49 +0.73 +0.53 +0.21 +0.38 +0.75 +0.55 +0.52 +0.39 +0.62 +0.16 +0.46 +0.63 +0.49 +0.46 +0.25 +0.16 +0.91 +0.71 +0.61 +0.36 +0.64 +0.51 +0.87 +0.38 +0.64

40

Ar(r) (%)

43.17 28.16 30.64 62.68 14.67 14.63 43.69 23.42 55.22 46.23 36.46 33.64 38.98 44.40 57.69 63.30 45.33 56.01 21.48 50.22 32.29 10.88 41.60 70.46 41.70 52.81 36.91 50.52 56.30 58.55 27.75 44.28 44.04 43.77 39.79 13.16 55.84 47.08 59.64 64.99 42.41 56.64 45.98 49.53 52.18 70.39 52.18 49.86 43.90 71.14 48.50 57.39 55.76 58.96 65.04 50.78 54.76 55.42 65.22

39

Ar(k) (%)

0.51 0.36 1.29 0.52 1.98 0.67 1.21 0.49 1.83 0.44 1.18 0.37 0.30 0.41 0.64 0.52 1.98 0.35 2.43 0.49 1.62 1.43 0.92 0.56 0.51 0.74 0.93 0.92 1.78 0.40 2.10 0.66 1.44 0.63 0.37 3.40 2.33 1.13 0.34 0.75 0.63 0.75 0.43 3.90 0.88 0.47 0.91 1.00 2.30 2.51 0.32 0.53 0.44 1.11 0.44 1.13 0.28 1.25 0.62 (Continued)

198

P. D. CLIFT ET AL.

Table 6. Continued TH-4-6 Age (Ma) 7.47 7.82 8.21 8.27 8.37 8.39 8.62 8.87 9.42 11.29 11.61 12.73 14.18 16.09 16.63 17.25 17.86 18.13 18.67 19.12 19.86 20.08 21.91 25.48 27.75 29.01 31.37 41.00 53.50 92.92 111.32

+2s +0.69 +1.03 +0.59 +0.51 +0.69 +0.44 +0.47 +0.57 +0.17 +0.27 +0.30 +0.94 +0.21 +0.38 +0.34 +0.25 +0.80 +0.25 +0.54 +0.49 +0.57 +0.71 +0.32 +1.34 +0.25 +0.50 +0.45 +0.63 +0.71 +0.70 +0.62

40

Ar(r) (%)

41.98 61.43 40.77 24.57 60.82 75.78 52.41 61.82 62.58 56.83 65.97 80.57 68.42 85.05 79.19 81.98 91.50 75.91 80.59 85.30 71.92 82.82 79.92 75.56 87.03 75.20 76.57 76.73 92.32 95.18 97.04

39

Ar(k) (%)

0.62 0.25 0.62 2.36 0.61 0.58 0.92 0.42 2.73 2.71 0.97 0.27 3.69 1.36 1.26 1.60 0.31 2.37 0.58 0.84 0.49 0.41 1.36 0.23 1.94 1.16 1.12 0.82 0.64 0.57 1.26

Canberra, using Excimer Laser Ablation Inductively Coupled Plasma Mass Spectrometry (ELA-ICP-MS) employing a pulsed 193 mm ArF LambdaPhysik LPX 1201 UV Excimer laser and an Agilent 7500 quadrupole ICP-MS. The zircons were separated from the bulk sediment by conventional magnetic and heavy liquid separation techniques. In this study we analysed two samples (Jati-16-1 and KB-34-1), which were not considered by the earlier study of Clift et al. (2008). The extracted zircons were mounted in epoxy resin and polished. Dating by ELA-ICP-MS followed the procedure described in Campbell et al. (2005). Our method employs standard zircon TEMORA2 and NIST610 silicate glass (Pearce et al. 1997; Black et al. 2004) where the latter is used for concentration information and for U/Th determination. As we cannot measure common Pb (204Pb) directly because of systemic Hg, we use a 208Pb-based correction only when that correction makes the analysis more concordant than the uncorrected version. Once the data were compiled, an analysis was rejected for

interpretation on the basis of the following: (a) the observed variance on 206Pb/238U or 207Pb/206Pb (depending if the grain is . or ,1200 Ma) is more than three times that calculated from counting statistics (this procedure omits grains that record mixed ages), or (b) the grain is deemed to be discordant. Analysis time drift corrections were applied to both analytical sessions. Results of the U – Pb dating are shown in the supplementary material. Overall uncertainty on an individual measurement is about 1–2%.

Results Mineralogy Detrital modes in the Indus Holocene sands are typical of sediments derived from collision orogens (Garzanti & Ando` 2007), with medium quartz content and equally abundant feldspars and lithic grains (quartz 49 + 4%, feldspar 26 + 3%, lithics 25 + 3%). Using a Dickinson ternary diagram (Fig. 3a) the sands mostly plot within the ‘recycled orogen’ field, with minor overlap in the ‘dissected arc’ range. In this respect they are similar to sandstones found in the Himalayan foreland basin, at least since the start of the Neogene (Najman & Garzanti 2000). Lithic grains comprise equally abundant sedimentary (limestone, dolostone, siltstone) and metamorphic types. However, volcanic, metavolcanic, and metabasite lithic grains are very minor, but decrease up-section in the lower part of the core (120 –60 m depth, 20 –9.5 ka; Fig. 4), with the greatest change between 12.2 and 10.6 ka. This trend suggests upward-decreasing relative contributions from Kohistan-like Transhimalayan sources and/or the West Pakistan ophiolites. The composition of the lithic grains defines the sands as typical of ‘suture belts’ in the ternary diagram (Fig. 3b). Conversely, carbonate grains, which are almost absent at the base of the Keti Bandar section (LGM) increase in the uppermost 90 m (since c. 12 ka), suggesting increasing contributions from sedimentary to metasedimentary strata and/or more arid climatic conditions that do not favour dissolution of these grains. The total heavy-mineral content is remarkably constant (4.2 + 0.7%) in the six very fine-grained sands analysed (KB-5-2, TH-10-8, KB-20-1, KB-34-4, KB-40-5, and KB-41-2), indicating that intrastratal solution is negligible throughout the cored sections. Heavy minerals are less abundant in the analysed 63 –250 mm fraction of the other three silt-sized samples (KB-23-2, KB-26-2, and KB-30-4), but this does not mean that bulk samples contain fewer heavy minerals, because denser detrital grains are markedly concentrated in the fine tail of the size distribution of each sample

MONSOON CONTROL OVER EROSION

(a)

Recycled Orogen

Ba se me n

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s. c o

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Transitional arc Undissected arc

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199

(ZTR) 3 + 1% and all other heavy minerals 6 + 2%) (Garzanti et al. 2005). Slightly higher garnet content in modern bedload sand is ascribed to the selective entrainment of less dense grains and enrichment in denser grains in lag deposits on the channel bottom (Slingerland 1984). Most remarkable is the virtual lack of limestone grains and of pyroxene at the bottom of the Keti Bandar core (117–120 m depth; i.e. deposited at the LGM), which would indicate strong chemical weathering. This observation may be linked to a humid climate at the LGM, for which there is no evidence, or alternatively and more likely, due to prolonged exposure prior to the Holocene transgression.

Lm

Sr isotope evolution Suture belts

Mixed arc and subduction complex

Magmatic arc

Lv

Ls

Fig. 3. (a) Triangular QFL plot (Q, quartz, L, lithics, F, feldspar) and (b) Lm (lithic metamorphic), Lv (lithic volcanic) Ls (lithic sedimentary) plots, with fields from Dickinson (1985) for the Indus Holocene sands.

because of settling-equivalence effects (Garzanti et al. 2009). Amphiboles (mainly blue-green hornblende, comprising 52 + 5% of the total heavy mineral population) prevail over epidote (27 + 4%), subordinate garnet (6 + 4%) and clinopyroxene (6 + 3%), and minor tourmaline, titanite, hypersthene, kyanite, sillimanite, rutile, staurolite, chloritoid and zircon in order of decreasing abundance (5 + 3%). Hypersthene increases slightly up-section, suggesting increasing contribution from the Kohistan Arc in the upper part of the core (depth ,60 m; since c. 9 ka), since this mineral is distinctive of erosion from arc units (Cerveny et al. 1989). Marked temporal changes in mineralogy are not apparent (Fig. 4), and sediment composition remains comparable to the modern Indus detrital modes (quartz 48 + 4%, feldspar 21 + 3%, lithics 32 + 5%) and heavy-mineral assemblages (amphibole 50 + 8%, epidote 25 + 7%, garnet 12 + 3%, pyroxene 4 + 3%, zircon þ tourmaline þ rutile

Temporal evolution in Sr isotopes is shown together with changing Nd isotopes in Figure 5. What is apparent is that the two systems are closely correlated and that Sr increases rapidly from low values at the LGM and until c. 13 ka, after which there is a rapid rise to a high 87Sr/86Sr between 8 and 9 ka. Subsequently, 87Sr/86Sr values decrease slightly to the present day. The correlation suggests that Sr isotope composition is dominated by source and provenance rather than by chemical weathering intensity, although the total number of analyses is rather low. As with the Nd data there is no clear correlation between isotopic composition and grain size. Sediments of the same grain size show a variety of isotopic ratios (Fig. 5). The shift to higher isotope ratios is consistent with increased relative erosion from the radiogenic crust of the Lesser and Greater Himalaya, and away from the more primitive crust of Kohistan and the Transhimalaya (Trivedi et al. 1984; France-Lanord & Le Fort 1988; Scaillet et al. 1990; Ahmad et al. 2000). Because the change in isotope character occurs at a time of greatly increased sediment flux to the delta we rule out the possibility that the change in isotope character reflects decreased sediment flux from the arc sources lying north of the Himalayan ranges and in any case the increase in hypersthenes (but not metabasic rocks) up-section suggests more erosion from Kohistan, not less. However, the trend in the Nd isotope curve indicates that the influence of Kohistan on the total sediment composition is swamped by increased flux from Himalayan sources.

Organic carbon Although the shift to higher 87Sr/86Sr values in the Early Holocene cannot be interpreted as indicating stronger chemical weathering under the influence of a stronger summer monsoon, organic carbon

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Proportion of total mineralogy (%)

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Proportion of heavy mineral population (%)

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50

Amphibole Group Pyroxene Group Epidote Group Zircon, tourmaline,Ti-oxides, titanite, apatite,monazite Chloritoid, staurolite, andalusite, kyanite, sillimanite Garnet

Fig. 4. Plots showing the evolution in sand mineralogy at the Keti Bandar borehole since 20 ka. (a) Major mineral groups, and (b) only heavy minerals. Data are plotted from Tables 1 and 2. The plots confirm the overall lack of strong changes in mineralogy throughout the deglaciation process.

analysis can be used to constrain environmental conditions. Total organic carbon (TOC) concentrations are low and range between 0.16% and 0.44% (Table 4), consistent with the generally arid conditions in the Indus basin. There is no obvious TOC variation with age of sedimentation and most of the samples remain in a narrow range around 0.3% (Fig. 6a). Bulk organic carbon d13C varies between – 23.3‰ and –19.7‰, but only two samples with very low TOC (0.16% and 0.18%) have d13C lower than –21.7‰ (Fig. 6b). These

negative d13C values ( –22.6 and –23.3‰) are associated with the lowest TOC and correspond to sandy sediments, as indicated by low Al/Si ratios. In these two quartz-rich sands, rock-derived organic carbon is likely to be a major component of the total organic carbon content (Galy et al. 2008). Their stable isotopic composition is therefore not representative of modern organic carbon directly derived from the biosphere. Excluding these two samples, d13C shows a c. 2‰ variation around an average value of –20.6‰. In detail,

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Fig. 5. Diagram showing the variability in bulk sedimentary Nd and Sr isotope character since the Last Glacial Maximum. Sediments are from the Indus delta and the Indus Canyon. Letters indicate grain size of sediment: C, clay; SI, silt; FS, fine sand; S, sand. Black arrows indicating age control points and Nd data are from Clift et al. (2008). Sr data are from Table 3.

d13C reaches a minimum after the LGM and likely until the Younger Dryas. d13C increases between 12 ka and 5 ka and finally decreases again up to present (Fig. 6b). Organic carbon (OC) can be derived from a mixture of different sources: terrestrial OC derived from vegetation, soils and autotrophic production in the river and marine organic carbon. Historically, the Indus River is characterized by high sediment concentration, which limits autotrophic productivity

(Ittekkot & Arain 1986). Terrestrial inputs may therefore be derived from the vegetation present in the basin, either directly (plant debris) or indirectly (soil organic carbon). C3 plants have a considerable range in d13C. Arid ecosystems are enriched in 13C (as high as 222‰) (Farquhar et al. 1989), but closed canopy flora are depleted in 13C, with d13C values as low as 235‰ (van der Merwe & Medina 1989). However, the average C3 value is about – 26‰. In contrast, C4 plants have a much more

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(a)

(b)

Total Organic Carbon (%) 0.15

0.25

0.35

–23 0

5

5

10

10

–22

–21

–20

Age (ka)

0

d13C (‰)

15

15

LGM

LGM

Fig. 6. Diagrams showing (a) the evolution in total organic carbon (TOC) and (b) carbon isotope character of bulk sediments from the Indus delta. See Table 4. TOC values are generally small and show little variation with time. Lowest TOC values correspond to negative d13C values associated with reworking of old organic matter. The 2‰ positive shift between 12 ka and 5 ka is significant and may be related to an increase of C4 plants input. See Table 4 for data.

restricted d13C range, with an average d13C value around –13‰ (Deines 1980; Hattersley 1982; Collister et al. 1994). In the Arabian Sea, modern marine plankton has d13C values around –20‰ (Fontugne & Duplessy 1978) and its isotopic composition has likely remained fairly stable during the last 20 ka. The sedimentological study of this record clearly indicates a marine transgression during the Holocene. Therefore, the middle section of the record at Keti Bandar (15 –70 m) may have been influenced by marine organic carbon, whereas it might be expected to be negligible in the upper and lower sections. In these two terrestrial sections of the record, bulk organic carbon d13C indicates mixed C3/C4 vegetation in the Indus basin. The lowest part of the record (LGM to 12 ka) shows a slight decrease of the bulk organic carbon d13C from –20‰ to –21‰. This shift may be related to an increase of the C3 plants proportion in the basin during the time of deglaciation. The c. 2‰ positive shift between 12.9 ka and 4.1 ka probably reflects

increasing contribution of marine OC, although it might also indicate an increase in the proportion of C4 plants in the basin. Conversely, the reversed trend from 4.1 ka to present likely reflects decreasing contribution of marine organic carbon, or an increase of the C3 plants proportion in the basin. Comparison with published organic carbon data from the Bengal Fan system indicates that the Indus is relatively low in TOC (Galy et al. 2007b). Maximum values range up to 0.44% compared to .1.1%. However, like the Bengal sediment our data show a rough first order correlation between Al/Si ratios and TOC (Fig. 7). This is consistent with a control of the organic carbon content by the sediment properties, specifically a preferential association of organic carbon with fine grain sediments enriched in phyllosilicates (clays and micas) (Galy et al. 2008). Figure 7 shows that the slope on the TOC v. Al/Si chart, which characterize the organic carbon loading, is lower in the Indus basin than in the Bengal Fan, reflecting its more arid environmental conditions.

MONSOON CONTROL OVER EROSION

although with only 19 grains this sample may not be representative (Clift et al. 2004). The Greater and Lesser Himalaya, together with the Karakoram have yielded abundant young AFT ages that would be consistent with a source in those ranges. The slightly older ages seen in the reworked Miocene foreland sedimentary rocks of the Siwaliks argues against them being important since 8.7 ka, although they may partly be responsible for the older age population seen at the LGM. Other possible sources for the older grains deposited at the LGM are the Transhimalaya or Kohistan. Further source characterization is possible using Ar –Ar and U – Pb methods.

al Fan

0.35

Beng

De

lta

0.25

Ind

us

TOC (%)

0.30

0.20

0.15 0.00

0.10

0.20

0.30

203

0.40

0.50

Al/Si Fig. 7. Diagram showing the relationships between mud Al/Si and TOC for the Indus Holocene. Trend shows much lower slope than that recognized for the Bengal Fan (Galy et al. 2007b), suggesting much lower organic productivity.

Fission track analyses The results of the fission track analyses are shown graphically in Figure 8 in the form of radial plots that show the ages and uncertainties of single grain apatite and zircon grains (Galbraith 1990). Statistical analysis allows a central age to be assigned for each sample, with greater confidence for those samples with more abundant grains. In a complex system like the Indus 100 grains are needed for a robust result (Ruhl & Hodges 2005) and in several of these sands the numbers are so low that they are not useful. Three samples can be used to look at the general development in sediment source since the LGM. Sample KB-40-5 has a central age of 9.0 + 1 Ma, yet by 8.7 ka sample KB-19-4 shows a central age of only 4.4 + 1 Ma. A similar age is yielded by sample TH-10-8, deposited around 7 ka. The radial plots show that there is a minority population dating .10 Ma, but that in the younger sands in particular this is very minor. The change in apatite fission track (AFT) ages during the Holocene must reflect a change in provenance as the duration is not long enough for this to represent a change in source exhumation rates. Comparison of the age spectrum in the sediments with AFT ages from possible source terrains allows the changing erosion patterns to be constrained (Fig. 9). Probability density diagrams emphasize the young AFT ages of the younger sediment and show the ‘tail’ of grains older than 10 Ma seen in the LGM sediment but not since that time. Comparison with the modern Indus sediments shows a similar pattern to the recent sediments,

Mica ages Ar –Ar cooling ages in biotite and muscovite micas document the age that these grains cooled below c. 280 8C and 350 8C respectively (Hodges 2003). As exhumation is diachronous across the Himalaya these ages can be used as powerful provenance tools in modern and ancient South Asian sediment (White et al. 2002; Clift et al. 2004). Figure 10 shows the range of biotite cooling ages for two core samples and one modern river sand sample. The age spectra for the modern and 6.4 ka sand differ in one key aspect from the LGM sand at Keti Bandar, in the abundance of grains ,10 Ma. All samples show minority populations with older, albeit Cenozoic cooling ages, At the LGM, the most common age lies around 9 Ma, compared to c. 4 Ma for the younger sediments. Comparison with the rather limited number of bedrock analyses suggests that the Karakoram makes a relatively good match as a possible source at the LGM, although there are no published data for the Transhimalaya or Lesser Himalaya. Cooling ages in the Greater Himalaya largely range 10–24 Ma and do not account for the up-surge of young ages in the Holocene delta, although there is a significant population of 10–14 Ma grains in the modern river that may be derived from this area. Nanga Parbat is a possible source of the 1– 7 Ma grains seen in both younger samples. Additional source constraints are possible using the muscovite Ar –Ar ages reported by Clift et al. (2008). This system has the advantage over biotite in being more widely measured in the potential source regions. Figure 11 shows that like the biotite data the 6.4 ka and modern sediments have several muscovite grains dating ,10 Ma, which the LGM sediment does not contain. A probability maximum around 18 Ma at 6.4 ka and at the LGM correlates with known sources in the Lesser and Greater Himalaya, although there is some overlap between these sources that makes their separation hard with this method. However, we note that

204

P. D. CLIFT ET AL. 40 KB-40-5 (Apatite), >20 ka, LGM Central Age: 9±1 Ma P(X2): 0.0% Relative Error: 46% Number of grains: 70

KB-40-5 (Zircon), >20 ka, LGM Central Age: 35±4 Ma P(X2): 0.0% Relative Error: 65% Number of grains: 28

30 20

+2

200 150 100

10

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5

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+2 0 -2

% relative error

1 % relative error

70

28

7 0

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30

40

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10 12

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Precision (1/sigma)

Precision (1/sigma)

KB-5-2 (Apatite), 210 a depositional age Central Age: 6±1 Ma P(X2): 5% Relative Error: 40% Number of grains: 16

20 16 12

+2

8

0

6

40 30

KB-19-4 (Apatite), 8.7 Ka depositional age Central Age: 4.4±1 Ma P(X2): 0.0% Relative Error: 124% Number of grains: 44

20

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-2

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-2

1 % relative error

55 0

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% relative error

89

11

8

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Precision (1/sigma)

Precision (1/sigma) KB-23-3 (Apatite), 9.3 ka depositional age Central Age: 11±3 Ma P(X2): 3% Relative Error: 57% Number of grains: 16

40 30

20 16

TH-10-8 (Apatite), 7.0 ka depositional age Central Age: 4.4±1 Ma P(X2): 0.0% Relative Error: 80% Number of grains: 61

12 8

20

+2

5

+2 0

10 -2

5

0 -2

1 % relative error

76

1 13

0 10

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30

Precision (1/sigma)

% relative error

63

6

0 10

20

30

40

50

60

Precision (1/sigma)

Fig. 8. Radial plots (Galbraith 1990) showing the ages and uncertainties of single grain apatite and zircon grains within Holocene sands from the Indus delta. Locations of samples are shown in Figures 1 and 2.

Greater Himalayan muscovite ages peaks around 20 Ma, whereas the limited data from the Lesser Himalaya peak around 16 –18 Ma. The greatest probability peak in the detrital grains is younger

than 20 Ma, consistent with the Lesser Himalaya being the dominant source. Again the muscovite confirms that the Siwaliks are not dominant sediment sources because the cooling ages are generally

MONSOON CONTROL OVER EROSION

205

Fig. 9. Probability density plots showing the range of apatite fission track central ages for the sediment samples analysed versus the ages found in a variety of possible source terrains. Siwalik data is from Van der Beek et al. (2006). Karakoram data is from Zeitler (1985), Poupeau et al. (1991) and Foster et al. (1994). Greater Himalayan data are from Kumar et al. (1995), Sorkhabi et al. (1996), Searle et al. (1999), Jain et al. (2000), Thiede et al. (2004) and Bojar et al. (2005). Lesser Himalayan data are from Vannay et al. (2004) and Thiede et al. (2004). Pakistan Himalayan data are from Zeitler (1985). Transhimalayan data are from Zeitler (1985), Zeilinger et al. (2001), Clift et al. (2002a) and Kirstein et al. (2006). Figure shows preference to younger grain ages with the onset of the Holocene, consistent with relatively more erosion from the Greater and Lesser Himalaya and less from the Siwaliks and Transhimalaya. See Table 5 for data.

too young. As for the ,10 Ma muscovite grains bedrock data suggest either Nanga Parbat or the Lesser Himalaya as likely sources. A probability maximum at 15 –16 Ma in the modern river matches several known sources in the Lesser Himalaya but not Nanga Parbat. This population is less abundant in the 6.4 ka sample. We conclude that the mica dating argues for reduced erosion in the Karakoram and more erosion in the Lesser Himalaya or Nanga Parbat between the LGM and the Early Holocene. As these sources are both negative in 1Nd (Parrish & Hodges 1996; Whittington et al. 1999; Ahmad et al. 2000) increased relative flux in either could explain the observed bulk sediment Nd and Sr isotope evolution (Fig. 5).

Zircon dating Zircon U –Pb dating has proven an effective provenance tool in South Asia because it preserves the original age of crystallization of the source rocks, which varies significantly across the Himalayas and into Tibet (DeCelles et al. 2000). In this study we augment the data presented by Clift et al. (2008) with two additional samples in order to define the Holocene provenance evolution better. Figure 12 shows the age spectra for the four core samples, plus modern river data compared with

various source terrains. All the sediments show a large population with grains of ,150 Ma and a spread of other older grains. Some samples show particularly well developed groups. The LGM sands show many grains dated at 800–1100 Ma, reducing in number up-section. The modern river sand shows an especially large number of grains c. 1800 Ma compared to the older sediments. The new data are consistent with the older in suggesting significant erosion from the Karakoram, or the Transhimalaya, especially at the LGM. In addition, the zircon indicate increased erosion from the Lesser Himalaya going up section. Very few of the grains are young enough to match those measured from the Nanga Parbat gneiss. This resolves one of the ambiguities from the mica Ar –Ar data in separating the erosional flux from the Lesser Himalaya. Despite its dramatic exhumation history (Zeitler et al. 1993) it appears that Nanga Parbat is a modest contributor of sediment to the Indus. The provenance can be further quantifying by dividing up the detrital zircons into families. We choose 0–20 Ma grains to represent the flux from Nanga Parbat, 20 –55 Ma grains are rarely known outside the Karakoram Batholith. The range 55 –300 Ma is chosen as a suitable range for much of the activity in Kohistan and the Transhimalaya,

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Greater Himalaya

Nanga Parbat

Karakoram 0

10

20

30

40

50

60

Modern Indus at Thatta N = 95

0

10

20

30

40

50

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TH-4-6, Biotite, <6.4 ka N= 46

0

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30

40

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KB-41-2, Biotite, 25 ka N = 99

0

10

20

30

40

50

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Age (Ma) Fig. 10. Probability density plots showing the range of Ar–Ar cooling ages in biotite grains from three sand samples from the Indus delta. Top section shows known range of possible source ages from the Greater Himalaya (Copeland et al. 1990; Searle et al. 1992; Metcalfe 1993; Inger 1998; Stu¨we & Foster 2001; Godin et al. 2006; Wang et al. 2006), Nanga Parbat (Zeitler et al. 1989; Winslow et al. 1996; Treloar et al. 2000) and Karakoram (Searle et al. 1989; Brookfield & Reynolds 1990; Krol et al. 1996; Villa et al. 1996). See Table 6 for data.

300– 1400 Ma grains represent the Greater Himalaya and those older than 1400 Ma the Lesser Himalaya. Because of overlaps in age ranges such a budget is necessarily schematic but does show

the general trends in provenance evolution because the different ranges have preferred ages that are typical if not unique to them. In Figure 13 we plot pie charts to show how the bulk composition

MONSOON CONTROL OVER EROSION

207

Greater Himalaya

Probability

Nanga Parbat

Siwaliks

Lesser Himalaya Modern River at Thatta (48 grains)

Probability

0-10 Ma (N. Parb.+L. Him.) = 15% 10-40 Ma (L.+Gr.. Him.) = 80% >40 Ma (Transhimalaya) = 4%

TH-4-6, Muscovite, <6.4 ka (99 grains)

Probability

0-10 Ma (N. Parb.+L. Him.) = 25% 10-40 Ma (L.+Gr. Him.) = 70% >40 Ma (Transhimalaya) = 5%

KB-41-2, Muscovite, >20 ka (50 grains)

Probability

0-10 Ma (N. Parb.+L. Him.) = 4% 10-40 Ma (L.+Gr. Him.) = 84% >40 Ma (Transhimalaya) = 12%

0

10

20

30

40

Age (Ma) Fig. 11. Probability density plots showing the range of Ar–Ar cooling ages in muscovite grains in the glacial sample, at ,6.4 ka and in the modern river (2004), compared to those in possible source regions. Top section shows the known range of possible source ages from the Greater Himalaya within the Indus basin (Searle et al. 1992; Metcalfe 1993; Inger 1998; Walker et al. 1999, Stephenson et al. 2001), Lesser Himalaya (Catlos et al. 2001; Bollinger et al. 2004; Vannay et al. 2004), Nanga Parbat (Smith et al. 1992; George et al. 1995; Treloar et al. 2000), and Siwaliks (White et al. 2002; Szulc et al. 2006). Reprinted with permission from the Geological Society of America.

of the Indus sediments has changed from the LGM to the present day. Again the heavy influence of the Lesser Himalaya on the modern river is clear and contrasts even with the mid Holocene samples. Because of damming of the modern Indus, most notably at Tarbela (Fig. 1), some of

the anomaly in the modern sample may be anthropogenic, although dams do exist on many of the Himalaya tributaries too (e.g. the Mangla Dam on the Jhelum). All samples show very little sediment from Nanga Parbat sources, but a consistent dominant flux from the Greater Himalaya.

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Nanga Parbat Karakoram

Lesser Himalaya

Probability

Greater Himalaya Siwaliks

Probability

Thatta, Modern Indus (130 grains) 2% Nanga Parbat 19% Karakoram+Transhim. 38% Greater Himalaya. 41% Lesser Himalaya

Probability

TH-10-1, Age = 7.0 ka (186 grains) 0% Nanga Parbat 28% Karakoram+Transhim. 38% Greater Himalaya. 34% Lesser Himalaya

Jati-16-1, Age = 8.0 ka (74 grains) Probability

3% Nanga Parbat 44% Karakoram+Transhim. 31% Greater Himalaya. 22% Lesser Himalaya

Probability

KB-34-1, Age = 12.6 ka (117 grains) 2% Nanga Parbat 41% Karakoram+Transhim. 35% Greater Himalaya. 22% Lesser Himalaya

Probability

KB-40-1 and -41-2, Age > 20 ka (271 grains) 1% Nanga Parbat 40% Karakoram+Transhim. 42% Greater Himalaya. 16% Lesser Himalaya

0

1000

2000

3000

Age (Ma) Fig. 12. Probability density plots showing the range of U– Pb ages in detrital zircons compared with source terrain values. Top section shows known range of possible source ages from the Greater Himalaya (Gehrels et al. 2006), Karakoram (Le Fort et al. 1983; Parrish & Tirrul 1989; Scha¨rer et al. 1990; Fraser et al. 2001; Heuberger et al. 2007), Lesser Himalaya (Parrish & Hodges 1996; DeCelles et al. 2000; Chambers et al. 2008), Nanga Parbat (Zeitler & Chamberlain 1991; Zeitler et al. 1993), and the Siwaliks (DeCelles et al. 2000; Bernet et al. 2006). See supplementary material for data.

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Fig. 13. Pie diagrams showing the changes in the relative proportions of different zircon U –Pb age populations in sands from the Indus delta. See Figures 1 and 2 for sample locations. Population 0 –20 Ma is a proxy for flux from Nanga Parbat, whereas the 20– 55 Ma are likely from the Karakoram Batholith. 55– 300 Ma grains are dominantly from the Transhimalaya (Ladakh and Kohistan batholith). Grains dated 300–1400 Ma are typical of sources in the Greater Himalaya, while older grains are likely derived (directly or indirectly) from the Lesser Himalaya.

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P. D. CLIFT ET AL.

Table 7. Predicted percentages of eroded material from a variety of western Himalayan source in five Indus River sands spanning the LGM to present. The mean 1Nd values for each of the sources is a modal number derived from published Nd isotope measurements from the bedrock Source

1Nd

KB-40-1

KB-34-4

Jati-16

TH-10-8

TH-1

Depositional age (ka) Nanga Parbat % Karakoram and Transhimalaya % High Himalaya % Lesser Himalaya %

225 þ1 216 224

20 1 40 42 17

13 3 39 36 22

8 3 44 31 22

7 2 37 39 22

0 2 22 36 40

210.7 210.8

211.4 211.8

210.5 n/a

211.7 212.9

215.7 215.4

Predicted 1Nd Observed 1Nd

Source: Nanga Parbat data is from Clift et al. (2002b) and Whittington et al. (1999), Greater Himalaya data is from Ahmad et al. (2000), Deniel et al. (1987), Stern et al. (1989), France-Lanord et al. (1993), Parrish & Hodges (1996), Searle et al. (1997), Harrison et al. (1999), Whittington et al. (1999). Lesser Himalaya data is from Ahmad et al. (2000) and Parrish & Hodges (1996). Transhimalayan data is from Khan et al. (1997), Clift et al. (2000). Karakoram data is from Clift et al. (2002b) and Scha¨rer et al. (1990).

Monsoon and erosion patterns The different provenance proxies can be combined to generate a ‘best-fit’ sediment budget for the Indus since the LGM. Because of the large number of grains and samples and the good degree of separation between sources we choose to base the budget on the U –Pb zircon grains (Fig. 13), but then cross-check this by calculating what the Nd isotope composition of such a sediment mixture would be given the known range of Nd isotope characteristics in the sources. These predicted 1Nd values can then be compared with the actual measured values of these sediments (Clift et al. 2008). Because the Nd analyses are bulk analyses they would be expected to yield good averages of the sediment flux. In practice there are some significant departures between predicted and observed 1Nd values. This may reflect use of an inappropriate 1Nd value for the sources, yet we consider this unlikely because the modal values often lie close to measured values from the major modern rivers, which themselves should sample and average wide areas of the possible sources (Clift et al. 2002b). Alternatively, we suggest that the zircons, which are interpreted simply in the pie diagrams of Figure 13 are not as accurate as might be hoped in characterizing the total mass flux because of age overlap between the populations and sources. There exists a further possibility that there is shortterm variability in the sediment provenance that results in measurable differences in zircon populations for sediments that were deposited close together in time. In Table 7 we show a proposed erosion budget for the Indus based on five sand samples. We use the zircon populations shown in Figure 13 as a

starting model for estimating the flux from each of the major sediment sources, but we adjust the relative proportions from each source to provide a closer match with the measured 1Nd values. Because the flux from the Karakoram and the Transhimalaya are hard to resolve from one another we plot these together as a single source. In each case the percentage adjustment from the observed was not more than 3%, and usually +2% or less. Figure 14 shows how this synthesized mass flux varies with time. The evolutionary patterns defined by this synthesis budget reflect many of the trends seen in the single mineral plots. Greater Himalayan flux remains high, if slightly variable throughout the period, as might be expected. Erosion from Nanga Parbat and the Lesser Himalaya because stronger

Nanga Parbat Karakoram and Transhimalaya

Proportion of total sediment (%)

Discussion

Greater Himalaya Lesser Himalaya

50 40 30 20 10 0

0

5

10

15

20

Age (ka) Fig. 14. Plot showing the evolving flux in zircon populations during the Holocene. We highlight the fall in relative flux from the Karakoram and Transhimalaya, compared to a sharp rise in the Lesser Himalaya, especially since 8 ka.

MONSOON CONTROL OVER EROSION

after the LGM, with a further sharp rise between 7 ka and the present day. At the same time flux from the Karakoram and Transhimalaya suffered a major decline. These changes may in part be related to the damming of the trunk river at Tarbela, which would raise the relative flux from the eastern tributaries draining the Himalaya, although these too have been dammed. Nonetheless, this theory does not explain the shift to similar negative 1Nd values in the Early Holocene. In this case we infer a similar shift to greatly enhanced erosion of the Lesser Himalaya relative to the Karakoram, peaking around 9 ka. The Nd isotopes suggest a moderate fall in Lesser Himalayan erosion after that time and before the most recent increase in the past 250 years. The later change in zircon sources (after 7 ka) compared to the earlier changes in Nd isotopes (10–14 ka) may also reflect a real lag in the sediment transport process for the zircon crystals versus clay minerals. The large scale changes in provenance tracked by the Nd isotopes (Figs 5 and 15) reflect a shift from preferential erosion of terrains lying to the north, around the Indus Suture Zone at the LGM to more erosion of the frontal Lesser Himalayan ranges in the south by the early Holocene. We do not think that drainage reorganization is responsible for the changing sediment compositions, even though this process has been used to explain changes in Nd isotopes after around 5 Ma (Clift & Blusztajn 2005). However, in this case the shift to more negative 1Nd values between 12 and 8 ka would require gain of isotopically negative sources. The increase in sediment flux at that time rules out loss of sources with positive 1Nd values as an alternative explanation. There is no evidence that the Punjabi tributaries, the Ravi, Jellum, Chenab and Sutlej, were captured into the Indus as recently as this time. Indeed, studies of Holocene drainage on the eastern edge of the Indus catchment indicate loss of drainage from Himalayan sources (Ghose et al. 1979) since the LGM, which would drive the opposite provenance shift than that observed. We conclude that the changes are driven by changing rates of sediment supply, not the wholesale capture of the Punjabi tributaries. When we compare the observed change in erosional style with records of the SW monsoon then we see that intensification of the summer rains [as tracked by speleothems (Fleitmann et al. 2003; Sinha et al. 2005) and pollen assemblages (Herzschuh 2006)] correlates with the strong change in Sr and Nd between 12 and 9 ka. This is also the time of accelerated sediment flux to the delta. Because the sediment composition changes quickly and is quite different from the sand deposited at the LGM we can rule out the sedimentation pulse as being caused by enhanced transport of

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older glacially eroded sediment under the influence of the strong Early Holocene monsoon. In any case the Lesser Himalaya were not glaciated during the LGM (Owen & Benn 2005) and so the deglaciation process should not have directly affected the erosion of these ranges. Instead the stronger monsoon appears to be generating new sediment by erosion under its precipitation maximum. Satellite data shows that precipitation maxima in the western Himalaya are focused over the topographic breaks in the Lesser and Greater Himalaya (Bookhagen & Burbank 2006) and that the change in provenance was probably caused by a strengthening of rain and erosion in those zones. Studies of landslides in the western Himalayan confirms that the Early Holocene was a period of significant mass wasting and thus sediment production within the Himalaya (Bookhagen et al. 2005). In contrast, erosion in the Karakoram appears to be largely glacially driven and would have been strong at the LGM, as well as today. Because the Karakoram lie in the rain shadow of the Himalaya and derive most of their water via the westerly jet (Karim & Veizer 2002) their erosion would not change significantly as the summer monsoon intensified.

Tectonics and the monsoon The primary conclusion of this study is that patterns of erosion across the western Himalaya changed significantly since 20 ka, driven by changes in summer monsoon intensity. Controls on erosion are important when considering the tectonic evolution of the Greater Himalaya because focused erosion is considered to have been a key factor in allowing deep buried metamorphic rocks to be exposed at the surface. This is true whether channelflow (Beaumont et al. 2001; Hodges et al. 2001) or orogenic wedge models are employed (Hilley & Strecker 2004) to explain the origin of the Greater Himalaya. Although we recognize that exhumation of ultra-high pressures, such as the Tso Moriri eclogites along the Indus Suture Zone are not erosionally driven (de Sigoyer et al. 2004; Leech et al. 2005), a purely tectonic origin for the Greater Himalaya is not currently favoured. Thermochronological transects across the Himalayan front suggest that zones of heavier precipitation correlate with areas of faster exhumation (Thiede et al. 2004; Wobus et al. 2005), although some indicators have been used to argue that rock uplift rather than monsoon rains dominate as drivers of erosion (Burbank et al. 2003). Furthermore, climaticallyfocused erosion appears to guide the location of active faults along the Himalayan front (Wobus et al. 2003). Our study reinforces the role of the monsoon in controlling orogenic architecture in the Himalaya. Without a strong summer monsoon

212 P. D. CLIFT ET AL. Fig. 15. Diagram showing the variability in various sediment and environmental proxies since the Last Glacial Maximum. Sediments are from the Indus delta and the Indus Canyon. (b) Nd data are from Clift et al. (2008). (c) Sr data are from Table 3. The Nd and Sr record are compared with (a) the GISP2 ice core climate record (Stuiver & Grootes 2000), (d) the variations in organic carbon isotope composition and (e) the intensity of the SW monsoon traced by speleothem records from Qunf and Timta Caves (Fleitmann et al. 2003; Sinha et al. 2005) in Oman and by pollen (Herzschuh 2006) from across Asia (black line), and well as western Himalayan landslides (Bookhagen et al. 2005). Note rapid change from C3 to C4 flora in early Holocene.

MONSOON CONTROL OVER EROSION

erosion is preferentially located in the Karakoram, such as at the LGM. However, for the exhumation of the Greater Himalaya to occur in the ways recently proposed a summer monsoon is crucial because the focused erosion required by such models does not occur in its absence. Our work also has implications for erosion/ tectonic coupling on a smaller scale. We show that the Nanga Parbat metamorphic massif, located in the western Himalayan syntaxis is a much less impressive sediment producer than its eastern twin at Namche Barwe, Tibet. France-Lanord et al. (2006) and Stewart et al. (2008) used a combination of thermochronological and U – Pb zircon data to indicate that as much as 45% of the sediment in the Brahmaputra is derived from erosion of the Namche Barwe massif, representing only 2% of the drainage system. In contrast, Nanga Parbat does not seem to contribute more than c. 3% of the sediment reaching the Indus delta. Our data casts some doubt over the idea that erosional unroofing by the Indus is driving the exhumation of the deep buried rocks (Zeitler et al. 2001) and that tectonic exhumation processes might also be significant (Hubbard et al. 1995). Why the two syntaxes behave in such different ways is unclear, although it is noteworthy that the Indus basin is much drier and generally much less erosive than the eastern Himalaya. Even within that region the effect of monsoon rain strength is a primary control on erosion rates (Galy & France-Lanord 2001).

Monsoon and the environment The organic carbon data now for the first time allow us to examine the changing environments in the Indus drainage basin. Other climate indicators, such as lake sediments (Enzel et al. 1999) from the edge of the Thar Desert show that after the Early Holocene maximum summer monsoon strength declined towards the present day. The d13C record from Keti Bandar shows significant correlation with the speleothem climate records (Fig. 15). There is a minimum in d13C following the Younger Dryas at c. 12 ka during the earliest Holocene, and then a rapid rise in the Early Holocene. The low d13C values seen at 8.6 and 9.8 ka are interpreted to indicate a dominant input from rock-derived organic carbon, but even excluding these points there is a rise d13C values during the Early Holocene. We interpret the change in d13C to reflect an increase in marine organic carbon flux into the sediment at Keti Bandar, as result of the marine transgression. Interestingly, d13C stays high well after the speleothem records start to decline but then decreases rapidly after c. 4 ka, when the sedimentation again becomes fluvial. This indicates rapid reduction in

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marine organic material after 4 ka. The carbon isotope data suggest that the summer monsoons in western India and Pakistan may not perfectly track those affecting Arabia and recorded in the Oman speleothem records (Fleitmann et al. 2003). The change in carbon isotopes has occurred most dramatically in the last 4 ka, and does not decrease gradually from 8 ka, as seen in the speleothems. A more detailed study, especially one using biomarkers, would be required to determine in detail when floral changes occurred, yet it is noteworthy that lake records from the environs of the Thar Desert show the most intense drying after 4.2 ka and were preceded by a much wetter period prior to that (Enzel et al. 1999).

Conclusions In this study we employed a series of petrographic, geochemical and isotopic methods to examine the effect on climate change since 20 ka on the nature of erosion and environmental conditions in the Indus river basin. Although the mineralogy of the sediments did not change much during this time there are coherent changes towards more radiogenic Sr and unradiogenic Nd isotopes that reflect increasing erosion of ancient crust between 12 and 8 ka. This represents the transition from Younger Dryas to Early Holocene and is known as a period when summer monsoon rains strengthened (Fleitmann et al. 2003). AFT data shows that since the LGM the Indus preferentially eroded sources that have younger ages compared to those at 20 ka (more rapidly eroded sources). Ar–Ar mica dates also show a shift to younger cooling ages at this time and together with U –Pb zircon dating demonstrates that the greatest change has been a relative decrease in erosion from north of the Indus Suture (i.e. from the Karakoram and Transhimalaya) and an increase in erosion from the Lesser Himalaya. As these ranges now lie in the zone of heaviest monsoon rains (Bookhagen & Burbank 2006) we infer that the change in provenance deposited at the delta is caused by erosion modulated by monsoon intensity. We can rule out remobilization of glacially eroded sediments. The close correlation with the climate history also indicates that sediment flux from source to delta was rapid and lower than the uncertainties in the 14C dating. Carbon isotope data also argue for a change in environmental conditions, with a sharp change to more positive marine organic carbon values during the start of the Holocene, as sea-level rose. Minimum 1Nd values are followed by a moderate increase after 9 ka, before a further decrease in the past 250 years. We interpret this to reflect damming of the main Indus at Tarbela blocking the flux of

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sediment from the suture zone. A drop in d13C values starting around 4 ka indicates a shift towards more terrestrial sources of organic carbon in the Indus basin, as the coast prograded southwards. The sediment record in the delta shows that erosion of the western Himalaya is strongly regulated by monsoon climate. Because current tectonic models for the exhumation of high grade metamorphic rocks in the Greater Himalaya require focused erosion our study suggests that formation of that range is dependent not just on the presence of tectonically thickened crust under southern Tibet but also on the activity of a strong summer monsoon. The pre-Holocene record shows that without the summer rains erosion is focused within and north of the Indus Suture and that consequently Himalayan exhumation cannot occur, at least not in the ways presently favoured by structural geologists. We thank US National Science Foundation (Ocean Sciences) for support of this project.

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Large-scale erosional response of SE Asia to monsoon evolution reconstructed from sedimentary records of the Song Hong-Yinggehai and Qiongdongnan basins, South China Sea LONG VAN HOANG1*, PETER D. CLIFT1, ANNE M. SCHWAB2, MADS HUUSE1, DUC ANH NGUYEN3 & SUN ZHEN4 1

Department of Geology & Petroleum Geology, School of Physical Sciences, University of Aberdeen, Meston Building, Aberdeen AB24 3UE, UK

2

Marathon International Petroleum (GB) Ltd, Marathon House, Rubislaw Hill, Anderson Drive, Aberdeen AB15 6FZ, UK 3

Vietnam Petroleum Institute, Yen Hoa, Cau Giay, Hanoi, Vietnam

4

South China Sea Institute of Oceanography, Chinese Academy of Sciences, 164 Xingang Road, Guangzhou, 510301 China *Corresponding author (e-mail: [email protected])

Abstract: The Song Hong-Yinggehai (SH-Y) and Qiongdongnan (Qi) basins together form one of the largest Cenozoic sedimentary basins in SE Asia. Here we present new records based on the analysis of seismic data, which we compare to geochemical data derived from cores from Ocean Drilling Program (ODP) Site 1148 in order to derive proxies for continental weathering and thus constrain summer monsoon intensity. The SH-Y Basin started opening during the Late Paleocene–Eocene. Two inversion phases are recognized to have occurred at c. 34 Ma and c. 15 Ma. The Qi Basin developed on the northern, rifted margin of South China Sea, within which a large canyon developed in a NE–SW direction. Geochemical and mineralogical data show that chemical weathering has gradually decreased in SE Asia after c. 25 Ma, whereas physical erosion became stronger, especially after c. 12 Ma. Summer monsoon intensification drove periods of faster erosion after 3–4 Ma and from 10– 15 Ma, although the initial pulse of eroded sediment at 29.5– 21 Ma was probably triggered by tectonic uplift because this precedes monsoon intensification at c. 22 Ma. Clay mineralogy indicates more physical erosion together with high sedimentation rates after c. 12 Ma suggesting a period of strong summer monsoon in the Mid-Miocene.

The history and causes for the East Asia Monsoon are controversial topics, although its proposed links with the uplift of the Tibetan Plateau remain a classic example of how the solid Earth may control atmospheric processes and the climatic evolution of the planet (Prell & Kutzbach 1992; Molnar et al. 1993). Certainly the presence of a large modern plateau plays an important role in driving the present intensity of the summer monsoon (Manabe & Terpstra 1974; Prell & Kutzbach 1992). In contrast, the winter monsoon is characterized by a cold and dry climate caused by air circulation in the reverse direction, although this too is linked with the growth of topography. As well as affecting atmospheric circulation patterns, uplift and deformation of the Tibetan Plateau has also intensified chemical weathering and physical erosion of source rocks as a result of

changes in rock physical properties and increasing terrain gradient. This in turn has affected the composition of sediments washed to the oceans by the large rivers that drain the eastern flank of the plateau. In this paper we assess the role of the Asian monsoon in controlling continental erosion. Although precipitation has been recognized as an important control on erosion (Reiners et al. 2003; Wobus et al. 2003) its relative role compared to tectonically driven rock uplift is unclear (Burbank et al. 2003). The South China Sea is a good place to examine the competing effects of these processes because the nature of the monsoon has been partly reconstructed from studies at a series of Ocean Drilling Program (ODP) sites on the rifted southern margin of China (Chen et al. 2003; Jia et al. 2003; Wan et al. 2006, 2007; Clift et al. 2008c). The

From: Clift, P. D., Tada, R. & Zheng, H. (eds) Monsoon Evolution and Tectonics –Climate Linkage in Asia. Geological Society, London, Special Publications, 342, 219–244. DOI: 10.1144/SP342.13 0305-8719/10/$15.00 # The Geological Society of London 2010.

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weathering and erosion records we present here can be readily compared with monsoon intensities to assess possible linkages. Although studies of the modern Red River suggest a dominant role for tectonically driven rock uplift in driving erosion (Clift et al. 2006b) this may not be the case over longer periods of geological time.

Location and geological setting Sediments eroded from eastern Tibet have partly been fed into the SH-Y and Qi basins via the Red River (Fig. 1). Although the Red River is still a large river it has been argued that the present drainage reflects major re-organization caused by re-tilting of eastern Asia towards the east during the Cenozoic (Wang 2004). Prior to this tilting the Red River may have formed the dominant drainage in SE Asia, but would have progressively lost drainage area because of headwater capture into adjacent systems (Brookfield 1998; Clark et al. 2004). Alternative models propose a more stable drainage and explain the curious geometries of river in SE Tibet as reflecting deformation, with

the rivers acting as passive strain markers (Hallet & Molnar 2001). Mass balancing eroded and deposited volumes of sediments now present in the SH-Y Basin and onshore in the modern Red River drainage indicated that the original catchment area of the Red River must have been much larger than that observed today (Clift et al. 2006a). Furthermore, Nd isotope values of sediments from the Hanoi Trough, Vietnam (Fig. 1) show a rapid change during the Oligocene (Clift et al. 2006a). Clift et al. (2006a) interpreted these changes as a response to large-scale drainage capture away from the former Red River. Hainan island is not considered to have been a major sediment source until it was uplifted during a period of strong magmatism that started c. 2 Ma (Tu et al. 1991). The influence of local sources along the Vietnamese coast on the total sediment influx is relatively poorly known. The SH-Y and Qi basins together form one of the largest sedimentary systems in SE Asia and are mostly filled by sediment delivered by the Red River. Thus, the sedimentary successions preserved within them record the history of erosion onshore. Although tectonic work can constrain the nature

Fig. 1. Location of the research area relative to SE Asia. The SH-Y Basin lies along the SE extension of the RRFZ, while the Qi Basin is situated on the rifted margin of South China Sea. The black straight lines show the location and length of the 2D seismic survey lines that were newly released to this study, while the red lines show those from Clift & Sun (2006). Black-dash lines are seismic lines selected for decompaction and sediment budget estimation. The white circles show the locations of the industrial wells used for the age assignment in this study. The black circle marks the location of ODP Site 1148 (Wang et al. 2000). Locations of seismic profile figures are shown.

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of deformation in Tibet and palaeoceanographic studies can reconstruct climate, it is the sediment records in the river deltas and fans that allow us to quantify erosion and so test for any links between climate, tectonics and erosion. Understanding how changing monsoon strength affects terrestrial environments is important not just for scientific reasons, but also because almost two-thirds of the global population are influenced by the monsoonal climate (Clift & Plumb 2008). Determining the relative role of solar insolation and atmospheric chemistry compared to tectonic processes in governing monsoon strength is important to predictions of future monsoon variability. The SH-Y and Qi sedimentary basins are situated within the Gulf of Tonkin, in the northwestern South China Sea (Fig. 1). The SH-Y Basin lies along the southeastern extension of the strike-slip Ailao Shan-Red River Fault Zone (RRFZ); the Qi Basin is situated at the southwestern end of the northern, rifted margin of South China Sea. The basins lie in two different tectonic settings. There is general agreement that formation of the basins was linked to opening of the South China Sea and motion on the RRFZ; whether these are all linked remains controversial. Tapponnier et al. (1982) carried out analogue experiments suggesting that penetration of the rigid Indian Plate into a softer Eurasia led to the extrusion of Indochina to the SE along the left-lateral RRFZ and consequently to the opening of the South China Sea. However, others argue that the opening of the South China Sea was triggered by a subduction force to the south where the Dangerous Grounds underthrust the Borneo Trench (Holloway 1982; Hall 1996; Morley 2002; Clift et al. 2008a). The SH-Y is interpreted as a pull-apart basin developed in a NW –SE orientation and controlled by a series of transtensional faults, especially the RRFZ, whose main trace is located on the SW side of the basin and by the ‘No.1 Fault’ to the NE. Modelling has demonstrated that a moderate degree of strike-slip shear caused by the rotation of Indochina relative to mainland China is capable of forming the basin geometry observed by seismic methods (Sun et al. 2003), without the need for motion .1000 km as had been suggested (Briais et al. 1993; Replumaz & Tapponnier 2003). The timing of deformation within the basin was dated by seismic methods to be before 30 Ma and to have ceased by c. 5.5 Ma (Rangin et al. 1995). Subsequently, Harrison et al. (1996), Wang et al. (1998), Leloup et al. (2001) and Gilley et al. (2003) all used radiometric isotope data to constrain the start of motion on the RRFZ to being close to c. 34 –35 Ma, broadly consistent with an acceleration of tectonic subsidence in the basin at that time (Clift & Sun 2006).

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In contrast, the Qi Basin was formed by rifting of the continental margin followed by seafloor spreading that started c. 30 Ma, as dated by marine magnetic anomalies in the neighbouring oceanic crust dated back to anomaly 11 (c. 31 Ma) (Taylor & Hayes 1980; Lu et al. 1987; Briais et al. 1993; Zhou et al. 2002), although there is a suggestion that seafloor spreading may date back to 37 Ma in the NE parts of the basin (Hsu et al. 2005). In any case the Qi Basin partly overlies both the continental shelf and continental slope (south of Hainan island), where water depth varies from c. 200–1500 m.

Monsoon reconstructions We have some knowledge of how the East Asian summer monsoon varies as a result of studies based on sediment records from ODP sites in northern South China Sea (ODP Sites 1146 and 1148; Fig. 1). Zheng et al. (2004) studied the abundance and ratio of planktonic foraminifera, which is a common proxy for reconstructing palaeoclimate change. They proposed that a decrease of the ratio of planktonic foraminifera Globigerinoides sacculifer/G. ruber and increase of Neogloboquadrina at c. 8 Ma at ODP Site 1146 indicates a lowering of the surface temperature and increased productivity, which are interpreted to have been caused by intensified East Asian winter monsoon winds. Upwelling-related radiolarian palaeomonsoon proxies in the southern South China Sea suggest that the east Asian summer monsoon first initiated close to the middle/late Miocene boundary at c. 12– 11 Ma and reached a maximum strength at 8.8 –7.7 Ma (Chen et al. 2003). This suggestion is consistent with work by Wan et al. (2007) who used sediment grain-sizes at ODP Site 1146 to indicate a stronger winter monsoon at c. 8 Ma and both winter and summer monsoon intensification at c. 3 Ma. However, geochemical data derived from ODP Site 1148 show much earlier intense chemical weathering in SE Asia, which may be linked to monsoon enhancement (Li et al. 2003). Continental weathering intensity is largely controlled by moisture and temperature and thus might be expected to be linked to the intensity of the summer monsoon rains. Wei et al. (2006) used a combination of the traditional, major element based ‘Chemical Index of Alteration’ (CIA) (Nesbitt et al. 1980), together with other geochemical proxies such as Ca/Ti, Na/Ti, Al/Ti, Al/Na, Al/K and La/Sm ratios to suggest that the summer East Asian monsoon has affected South China since the Early Miocene. Curiously, this study suggested that summer monsoon rains have gradually decreased while winter monsoon strength has increased since that time.

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Most recently a colour spectral-based analysis of clay mineralogy at ODP Site 1148 has shown more complicated and significant mismatches with other data set (Clift et al. 2008c). This record indicates an initial intensification of weathering after 22 Ma, followed by a period of especially strong summer monsoon from 16– 10 Ma. The summer monsoon would appear to weaken into the Pliocene before experiencing moderate intensification since c. 4 Ma. Earlier monsoon reconstructions sometimes show inconsistent results. In some cases this reflects different timescales and resolution of data and/or grain-size distribution of sediments. Other apparent discrepancies may reflect the fact that the different proxies are measuring different things, for example, upwelling, weathering, wind strength, which in turn may track activity of either winter or summer monsoons. The spectral data of Clift et al. (2008c) are derived from clays, whereas CIA is bulk sediment analysis that can also be affected by grain size.

Data sources In this paper we present new constraints on the history of sediment flux from the Red River Basin based on interpretation of newly released 2D multichannel seismic profiles from the SH-Y and Qi basins, which allow us to augment earlier seismic stratigraphic studies of the basins (Rangin et al. 1995; Fang et al. 2000; He et al. 2002; Gong & Li 2004; Clift & Sun 2006). In particular, our data adds greatly to our understanding of the southwestern, Vietnamese parts of the SH-Y Basin compared to the previous regional synthesis of Clift & Sun (2006) (Fig. 1), which was focused much more on the eastern half of the basin. These new data allow us to estimate the flux of sediment into the whole basin through time more completely. We have constrained changing weathering regimes in the area using geochemical data from XRF whole corescanning of sediment from Ocean Drilling Program (ODP) Site 1148 in the northern South China Sea (Fig. 1). Unfortunately, no suitable core exists for the Red River offshore, so here we have exploited a core from the neighbouring Pearl River drainage. This has the additional advantage of imaging a river basin that is largely unaffected by Neogene tectonics or drainage reorganization (Clark et al. 2004), so that changes in weathering regime can be readily related to climate and thus to the monsoon.

Methodology Seismic data and workflow In this study, we analysed 2D multichannel seismic profiles provided by BP, PetroVietnam and the

Chinese National Offshore Oil Company (CNOOC). In total, 48 lines (c. 5500 km) from the SH-Y Basin and 12 lines (c. 750 km) for the Qi Basin (Fig. 1) were used to constrain the sedimentary evolution. These data augment earlier published data, largely from the Chinese sector. There have been several interpretations carried out by oil and gas companies for different parts of the SH-Y Basin. However, geological correlation across the basin has not previously been possible because this basin straddles the international boundary between Vietnam and China. In this study, we used seismic data, which cover the whole area of the SH-Y Basin (Fig. 1) in order to have better geological interpretation and correlation from the SW to the NE side, as well as towards the SE end, where it meets the Qi Basin (Fig. 1). Once navigation and SEG Y data (SEG Y file format is one of several standards developed by the Society of Exploration Geophysicists for storing geophysical data) were loaded onto a workstation running KingdomTM, software seismic sequence boundaries were picked, based on conventional termination types of seismic reflections (e.g. onlap, downlap, erosion) (Vail et al. 1977; Miall 1991). Age constraints derived from biostratigraphy in industrial wells, typically at the sub-epoch level of resolution were assigned to these horizons prior to time– depth conversion.

Geochemical analysis Geochemical data have been used as a proxy for constraining the intensity of chemical weathering, which is controlled by a number of processes including temperature (White et al. 1999) and moisture (Gabet et al. 2006). These in turn may be linked to monsoon strength (Derry & France-Lanord 1996b; Filippelli 1997). In this study, we used cores from ODP Site 1148, located on the deep-water slope offshore the Pearl River (Fig. 1) to examine variations in major element chemistry. Cores from this site were analysed at the Research Centre Ocean Margin (RCOM), University of Bremen, Germany by an X-Ray Fluorescence Core Scanner manufactured by AVAATECH. One half of each core was flattened and covered by plastic film before being positioned under the X-ray beam for scanning. The step-size for each measurement was set up at every 7 cm. However, this resolution was changed in accordance with lithological variation and to avoid fractures within the core. XRF scanning can obtain continuous data at much finer scales than is practical for individual sampling methods. These advantages are especially important for relatively long time series and especially for highresolution analyses on critical boundaries/intervals. XRF core scan data show a significantly higher

TECTONIC AND MONSOON EVOLUTION IN SE ASIA

signal-to-noise ratio and more consistent holeto-hole agreement than standard logs. Use of XRF core scan data as a tool in palaeoceanographic and stratigraphic studies is well defined and widely accepted (Ro¨hl & Abrams 2000; Tjallingii et al. 2007). Use of major element data to constrain terrestrial weathering intensity often involves use of the CIA which is an established weathering proxy (Nesbitt & Young 1982). However, CIA is dependent on having Al, Na, K and Ca concentration data and is susceptible to variation linked to sediment mineralogy and provenance evolution. Unfortunately, the scanner does not provide any Na concentration data. Even if CIA can be determined, sands yield lower CIA values compared to clays in the same drainage system (Clift et al. 2008b).

Results Tectonic and seismic stratigraphic evolution Because the SH-Y and Qi basins are situated in different tectonic provinces, they were interpreted independently. By picking horizons along termination surfaces, fourteen sedimentary packages were defined for the SH-Y Basin, and the sedimentary formations of the Qi Basin were sub-divided into nine packages. The age for each boundary surface was dated by nannofossil-based biostratigraphy provided by the operating company from each industrial well and/or by correlating stratigraphy across regional cross-sections. The cross-section shown in Figure 2 shows the general structure of the SH-Y Basin. It shows a classic pull-apart type basin with a relatively symmetrical shape around an axis developed in a NNW–SSE direction (Dooley & McClay 1997). Because of the limited seismic coverage and lack of drilling data from the basin centre, the morphology of the basement as well as the nature of the oldest sedimentary formations in the basin centre has not been well defined. These deepest formations were estimated as being Paleocene –Eocene, unconformably overlying pre-Cenozoic terrigenous, carbonate sedimentary rocks with minor volcanic rocks (Tran et al. 2003; Mai et al. 2005; Clift & Sun 2006). The basin started to subside after c. 50 Ma, presumably related to the regional crustal extension seen in other parts of the South China Sea (Su et al. 1989; Clift & Lin 2001). This phase was followed by rapid subsidence especially after c. 34 Ma when motion on the RRFZ started (Gilley et al. 2003) and the pull-apart basin developed. Active tectonic subsidence continued until the Late Oligocene –Early Miocene. The extensional faults and carbonate platforms of the pre-Cenozoic basement have created a complicated basement

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morphology and thus a highly laterally variable sediment distribution (Figs 3, 4 & 5). In Figure 2, most of the syn-rift deposits are observed to be in the basin centre, where they are displaced by numerous transtensional faults on both margins. Offsets of the Palaeogene– Lower Miocene formations across these faults suggest that they were reactivated several times. Active rifting was followed by strong thermal subsidence during the Miocene, gradually weakening, but then accelerating again after Miocene (c. 5.5 Ma), at least in the southern SH-Y Basin (Clift & Sun 2006). Low-angle shoreline trajectories observed in the post-Miocene sedimentary packages demonstrate that little accommodation space was created at that time. As a result, most of the sediments eroded from Hainan island during the Pliocene quickly prograded towards the basin centre with very little vertical aggradation (Fig. 6). During its evolutionary history, the SH-Y Basin has experienced at least two inversion phases. The first uplift period is quite localized and is interpreted to have been triggered by the onset of the motion on the RRFZ at c. 34 –35 Ma. In contrast, the later Mid-Miocene event (c. 15.5 Ma) probably correlates with the generation of the Deep Regional Unconformity in the South China Sea (Hazebroek & Tan 1993; Hutchison 1996; Matthews et al. 1997), the end of the motion on the RRFZ, and with the cessation of seafloor spreading. Evidence for uplift and basin inversion is provided by strong deformation and erosion signatures observed on seismic profiles, especially in the northern part of the SH-Y Basin (Figs 3 & 7). The basin inversion involved not only deformation of basin fill, but also thrust faulting with significant vertical offsets (c. 250 ms of two-way travel time; Fig. 7). Compressional stresses operating during inversion may have played an important role in remobilizing the overpressured, fine-grained sediments to form shale/mud diapir-like structures in the centre of the SH-Y Basin (Hao et al. 2002; Figs 2 & 8b). However, some of these structures penetrate the youngest sedimentary formations that postdate Middle Miocene inversion. This relationship suggests that the structures were initially formed by compressional deformation but may subsequently have been enhanced by sedimentary loading during rapid deposition of the overlying Pliocene –Recent sediments (Clift & Sun 2006). Although these structures could conventionally be interpreted as shale diapirs (Hao et al. 2002; Xie et al. 2003; Clift & Sun 2006) their shapes and the structure of surrounding and overlying seismic reflections do not resemble geometries associated with well documented diapirs elsewhere. A possible alternative interpretation in better agreement with seismic stratigraphic relations may

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Fig. 2. Seismic and interpreted cross-section of Line GPGT 93-223 through the SH-Y Basin, showing a pull-apart basin structure. Rifting has strongly disrupted the basement and displaced syn-rift formations with different offsets. A possible shale/mud diapir is intruded into younger layers as a result of the sediment remobilisation initially triggered by tectonic inversion. Alternatively this structure could represent a strike-slip fault zone. Line location is shown on Figure 1.

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Fig. 3. Seismic and interpreted cross-section of Line GPGT 93–200 running in a north–south direction through the SH-Y Basin and showing a complicated basement morphology and the increased deformation/erosion towards the north. Line location is shown on Figure 1.

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Fig. 4. Seismic and interpreted cross-section of Line GPGT 93–204 in the SH-Y Basin. The presence of a carbonate platform reflects more localized sediment distribution during the early stages of basin opening. The progradation configuration observed here suggests sediments spilled over to the SE during the Plio-Pleistocene. Line location is shown on Figure 1.

TECTONIC AND MONSOON EVOLUTION IN SE ASIA

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Fig. 5. Seismic and interpreted cross-section of Line GPGT 93–225 on the SW flank of the SH-Y Basin. The progradation towards the ENE suggests more sediments spilled over from the northern SH-Y Basin were delivered into the Qi Basin after c. 2 Ma. Line location is shown on Figure 1.

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Fig. 6. Cross-section of seismic profile C-65-75 through the Qi Basin. Basement is characterized by strong faulting, which formed a succession of graben-horst structures. Thin layers show that less sediment was delivered to this basin before c. 2 Ma, while steep shelf edge trajectory suggests rapid sediment influx after this time. A large-scale canyon, which incised down through older formations shows an imbalance between the sediment supply and slope stability. Line location is shown on Figure 1.

TECTONIC AND MONSOON EVOLUTION IN SE ASIA

Fig. 7. Cross-section interpreted from seismic profile GPGT 93– 201 in the northern SH-Y Basin. Strong erosion, deformation and thrust faulting are evidence for several basin inversions. Line location is shown on Figure 1.

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Fig. 8. Several internal seismic reflection architectures observed within the basin: (a) asymmetrical channel migration; (b) shallow faulting developed above a shale diapir/fault zone; (c) complex-oblique sigmoidal reflections in the progradational package; (d) high amplitude/low frequency of the seismic reflections observed at the bottom of the canyon was likely caused by the presence of a carbonate-cemented sandstone. The underlying reflections are pulled up due to the increase in seismic velocity of infill within the canyon; (e) lateral variation in seismic facies; (f) vertical variation in bedded seismic facies.

include combined dip- and strike-slip faulting to create triangular zones of poor seismic imaging overlain by only subtly disturbed, sub-horizontal reflections (Figs 2 & 8b). However, the present data density does not allow a confident interpretation of these structures and their origin is not discussed further here.

In contrast to the SH-Y Basin, the Qi Basin overlies the rifted margin of southern China and straddles the continent–ocean transition (Hao et al. 1998). Its shape was strongly influenced by the palaeogeomorphology of the rifted continental margin basement. Seismic characteristics and the interpreted cross-section in Figure 6 show the

TECTONIC AND MONSOON EVOLUTION IN SE ASIA

large-scale internal architecture of the Qi Basin. Accurate age control is hard to achieve across this section because of a lack of drilling in the deep water. The basement of the Qi Basin is strongly disrupted by a series of normal faults to form a classic graben-horst structure (Fig. 6). Rifting appears to have ceased by 21 Ma, a little earlier than in the SH-Y Basin, but a little younger than seen in the main depocentre east of Hainan island, the Pearl River Mouth Basin. Sediment influx into the Qi basin before the Pliocene was limited and likely derived from southern China and Hainan because older sediment from the Red River was accommodated in the SH-Y Basin. The total volume of sediments deposited in the Qi Basin is much smaller compared to those in the SH-Y Basin. After c. 2 Ma, sedimentation in the Qi Basin became notably faster. The increase in sediment supply during this period is indicated by steep shoreline trajectories and progradation patterns observed on seismic profiles (Figs 6 & 8c). The vertical stacking and horizontal progradation patterns observed in the seismic profiles demonstrate that sediment supply was faster than accommodation space creation at that time. Another striking feature observed in this basin is the presence of a large-scale canyon, which developed in a NE –SW direction, and which widens and deepens towards the SE (Fig. 6b). The maximum observed size of the canyon is estimated to reach c. 30 km wide and c. 1 km deep. The canyon formed after c. 2 Ma and incised older sediments dating back to c. 2.6–3.6 Ma. Slope gradients increase locally associated with the canyon. As a result, more coarse-grained sediments filled the head of canyon, as evidenced by faster seismic interval velocities compared to the surrounding area. This is illustrated by strong amplitude, low frequency of seismic reflections at the base of the canyon and by the ‘pull-up’ effect of the underlying reflections (Fig. 8d). The formation mechanism of this canyon is unclear but may result from relative base level fall as the area around Hainan was uplifted causing down-cutting and cannibalization of the slope stratigraphy.

Time – depth conversion Because well data are only available down to limited depths, which are much shallower than basement in the basin centre, we used stacking velocities derived from seismic processing in order to make a time-depth conversion and thus estimate the depths of stratigraphic and basement surfaces. Once the major stratigraphic surfaces were picked, the sections were converted from time to depth using stacking velocities derived from seismic data processing. We calculated the interval velocity

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for each layer by applying Dix’s equation (Dix 1955): sffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffi V22  T2  V12  T1 Vi ¼ T2  T1 where: Vi is interval velocity for each layer V1, V2 are stacking velocity values for the upper and lower layer boundaries T1, T2 are two-way travel time values down to the upper and lower layer boundaries. The layer thickness and the bottom depth are computed by following equations: Li ¼

Vi (T2  T1 ) 2

and

Di ¼ Di1 þ Li

where: L1 is layer thickness D1, Di-1: are the depths of the upper and lower layer surfaces. The depth of each boundary surface was used to construct isopach maps and to estimate the sediment budget using all the sections shown in Figure 1. Figure 9 shows contoured isopach maps generated for major stratigraphic units within the SH-Y and Qi basins. Seismic interpretation shows that the basin centre has a maximum depth of c. 22 km. The variation in morphology and depth of the basin through time demonstrates that the depocentre was relatively stationary until the Middle Miocene, after which it gradually migrated towards the SE, following strong basin inversion in the north. In contrast, the Qi Basin overlies both continental shelf and slope, and its deposition pattern is different from the SH-Y Basin. Because this basin forms part of a rifted passive margin and most of the sediments were delivered from Hainan island, sediments tended to accumulate quickly on the continental shelf and slope, but more slowly in the deep basin floor. As a result, total sediment thicknesses in the northwestern half of the basin (c. 8.5 km) are much higher than those on the opposite flank (c. 4 km) (Figs 6 & 9a). Within the Qi Basin, the depocentre has not just migrated towards the SW where it intersects with the SH-Y Basin, but also to the SE (Fig. 9a).

Sediment budget Decompaction of the depth-converted sedimentary layers allows us to restore them to their original volumes by correcting for the hydrostatic load of the overlying layers (Sclater & Christie 1980). This procedure was executed in 2D using the software FlexDecompTM v.1.0 (Kusznir et al. 1995).

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Fig. 9. Isopach maps of the SH-Y and Qi Basins generated for different time periods show that the depocentre has migrated through time: (a) Total sedimentary thickness of the basin; (b) Eocene– Oligocene; (c) Early Miocene; (d) Pliocene– Pleistocene. The varying sediment thickness shows that more sediment has been deposited in the northwestern half of the Qi Basin, under the present shelf and slope than the southern half of the basin.

TECTONIC AND MONSOON EVOLUTION IN SE ASIA

In this study, we selected eight seismic lines from the SH-Y Basin and seven lines from the Qi Basin for backstripping. These lines are distributed across the whole basin, covering wide stretches in order to ensure their representative character (Fig. 1). The input data consist of depth values for each stratigraphic surface, lithology and age constraints. Because no age control was defined for the oldest sedimentary layers, we assumed the basin started opening at c. 50 Ma. Each individual layer was then decompacted in reverse order of deposition by unloading the overlying packages. The total unloaded area of each stratigraphic unit was normalized to the total area of the whole decompacted sections within the basin in order to define a normalized coefficient of erosion for each individual stratum. The true volume of sediment deposited during any given time period across the whole basin was estimated by multiplying the basin volume by the normalization coefficient. Rates of sediment supply were then derived by dividing by the duration of sedimentation. The sedimentation rate was computed not only for both SH-Y and Qi basins respectively, but was also calculated for the combined SH-Y and Qi basin. Details of the sediment budget estimation are presented in Figure 10 and Table 1, where our results can be directly compared with those from the earlier work by Clift & Sun (2006) and Me´tivier et al. (1999).

Temporal evolution in mass flux From c. 50–29.5 Ma, sedimentation rates in the SH-Y Basin were modest, but increased for the period 29.5 –21 Ma (Fig. 10a). Sedimentation rates fell again between 21 and 15.5 Ma before rising to a higher level at 15.5 –10.5 Ma. Maximum values of sediment supply are calculated for the Plio-Pleistocene, following a period of lower sedimentation at 10.5–5.5 Ma. This general pattern is comparable with that estimated by Clift & Sun (2006), although they predicted peak sedimentation in the Middle, not the Early Miocene, as we do here. In contrast, sedimentation rates in the Qi Basin are quite different (Fig. 10b). Average sedimentation rates were generally low, except for a modest increase around 15 Ma. The sediment flux into the Qi Basin spiked rapidly only after c. 3 Ma. In this general form our reconstruction is close to that of Clift & Sun (2006), although we predict much higher peak values in the Pleistocene. Because the SH-Y Basin is much larger than the Qi Basin, the combined sedimentation rate reconstruction is similar to that of the SH-Y Basin (Fig. 10c), although with an accentuated pulse of sediment delivery since 3 Ma. Our reconstruction differs greatly from the predicted gradual rise in rates

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predicted by Me´tivier et al. (1999) since the Eocene, but mostly differs from the budget of Clift & Sun (2006) in emphasizing the increase in rates in the Early Miocene, with a reduced flux in the Middle Miocene.

Monsoon weathering reconstructions In order to assess if climate is controlling temporal variations in sediment flux to the ocean we require a detailed history of environmental conditions with which to compare our sediment budget. In this study we assume that sedimentation rates are a proxy for erosion rates in the Red River Basin onshore and that continental weathering intensity can be used as a proxy for summer monsoon rains. We use weathering records derived from the Pearl River Mouth Basin, largely from ODP Site 1148, where the sediments are derived from erosion of largely flat-lying southern China (Li et al. 2003). The general lack of tectonism or major drainage capture means that variations in sedimentary composition largely reflect climatically modulated chemical weathering intensity. In this study, we used published weathering data (Wei et al. 2006; Wan et al. 2007; Clift et al. 2008c) together with new chemical proxies calculated from the XRF scanning data to trace different aspects of the clastic flux. Figure 11 shows a variety of mineralogical and geochemical records, which do not show parallel development. This raises a question concerning the reliability of the scanning data, as well as the other records, which are not in agreement. Kido et al. (2006) pointed out that XRF signal intensity is reduced by the presence of water within sediment and that a thin water film between the sediment surface and the covering film may affect the reliability of the output data. The presence of water between the sediment and the film may be important but it would be equally so for all parts of the core and would not account for coherent longterm trends in the data, although it might explain some of the short duration spikes in values, especially as seen in Al. Progressive dewatering in deeper buried sediments could generate a long-term trend, although we do not see anything that might be suggestive of that in our data. The absorption effect of water is in the following order Al . Si . K . Ti. Consequently, decreasing water content downsection should increase the intensity of Al counts relative to Si, K, and Ti. This may cause a decrease in Al/Si ratios up-section. However, core description and carbonate contents (Clift 2006) show us that the clastic content was relatively stable and only increased after 6–7 Ma, synchronous with the change in Al/Si values. In contrast, porosity changes more progressively, and only shows a

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Fig. 10. Sediment budget estimate derived from this study (dark grey) and compared to the earlier work of Clift & Sun (2006) shown in cross-hatched patterns and Me´tivier et al. (1999) shown with horizontal lines for: (a) the SH-Y Basin; (b) the Qi Basin; (c) total combined SH-Y and Qi Basin.

trend to higher water contents at depths shallower than c. 90 m, equivalent to ages of c. 1.3 Ma (Shipboard Scientific Party 2000). We conclude that there is no correlation between porosity and Al/Si values. This indicates that clay content rather than porosity is the primary control on Al/Si values.

We further compare the geochemical weathering records with clay mineral assemblages at nearby ODP Site 1146 (Wan et al. 2007) because certain clay minerals, such as kaolinite and smectite are formed by chemical weathering processes, whereas others, such as illite and chlorite, are the products

Table 1. Sediment budget estimation for the Song Hong-Yinggehai and Qiongdongnam basins Stratum area/Seismic line (km2) 3555

C-58-79

Song Hong–Yinggehai Basin 0.0–2.0 52.8 110.9 2.0–5.5 128.8 237.9 5.5–10.5 36.0 50.6 10.5–15.5 91.2 87.8 15.5–21.0 89.0 38.9 21.0–29.5 168.9 54.0 29.5–50.0 75.7 157.7

GPGT 93-207

GPGT 93-211

GPGT 93-215

GPGT 93-219

GPGT 93-223

GPGT 93-225

59.7 294.7 72.4 144.5 65.1 398.2 673.3

85.0 309.3 104.4 228.0 180.5 1003.8 1280.9

109.2 274.9 151.4 344.2 232.2 878.3 802.2

124.4 237.8 278.5 337.7 212.0 538.3 518.6

143.2 262.7 166.5 262.4 145.2 483.4 684.7

183.8 168.9 223.2 231.5 83.9 261.8 278.0

Total area (km2)

Real volume (km3)

Mean sedimentation rate (km3/Ma)

Maximum sedimentation rate (km3/Ma)

Minimum sedimentation rate (km3/Ma)

868.9 1915.2 1082.9 1727.4 1046.9 3786.8 4471.1

32.1 70.7 40.0 63.8 38.6 139.8 165.0

16.0 20.2 8.0 12.8 7.0 16.4 8.1

19.2 24.2 9.6 15.3 8.4 19.7 9.7

12.8 16.2 6.4 10.2 5.6 13.2 6.4

1079.0 482.8 117.7 186.7 180.7 141.8 347.9 162.3 121.8 919.4

86.5 38.7 9.4 15.0 14.5 11.4 27.9 13.0 9.8 73.7

43.3 24.2 5.0 3.0 4.4 6.7 5.1 3.8 3.2 3.3

51.9 29.0 6.0 3.6 5.3 8.0 6.1 4.6 3.8 3.9

34.6 19.4 4.0 2.4 3.5 5.4 4.1 3.1 2.5 2.6

Stratum area/Seismic line (km2) C-57-79 C-35-69 Qiongdongnan Basin 0.0–2.0 39.6 2.0–3.6 53.4 3.6–5.5 30.5 5.5–10.5 8.0 10.5–13.8 41.2 13.8–15.5 25.2 15.5–21.0 22.2 21.0–24.4 26.9 24.4–27.5 8.0 27.5–50 76.0

33.1 37.6 21.5 5.7 25.4 13.2 13.1 18.7 7.4 71.2

44039

C-49-79

C-65-79

C-73-79

C-98-79

76.9 43.7 5.4 14.2 9.4 5.9 19.0 17.1 15.6 113.2

159.3 61.8 8.6 22.7 15.0 14.0 45.0 18.5 16.9 122.7

191.2 77.4 12.5 32.9 21.7 15.0 49.8 20.0 18.2 132.1

267.3 106.4 17.3 45.5 30.0 30.5 98.8 25.0 22.8 165.5

311.7 102.5 21.9 57.7 38.0 38.0 100.0 36.1 32.9 238.7

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Stratum age (Ma)

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Fig. 11. Geochemical data derived from ODP Site 1148 and showing the temporal variations in the intensity of chemical weathering through the Neogene, interpreted as a monsoon proxy in this study: (a) colour data (CRAT) (Clift et al. 2008c); (b) illite/smectite ratios for ODP Site 1146 (Wan et al. 2007); (c) chemical index of alteration (CIA) (Wei et al. 2006); (d) Al/Si ratio; (e) Ti/Ca ratio; (f) sediment budgets estimated for the whole SH-Y and Qi Basins.

TECTONIC AND MONSOON EVOLUTION IN SE ASIA

of physical weathering (Thiry 2000). As a result ratios such as illite/smectite can be used to indicate relative strengthening of physical v. chemical processes (Fig. 11b). The data from ODP Site 1146 suggest strengthening of physical weathering after 12 Ma and especially after c. 8 Ma. High-resolution scanning-based weather proxies can also be compared with whole rock XRF analyses of Wei et al. (2006), which allow the CIA to be calculated (Fig. 11c). This proxy shows a general trend to decreasing values through time, which Wei et al. (2006) interpreted to reflect decreasing weathering intensities under the influence of a steadily weakening summer monsoon. This simple interpretation is hard to reconcile with some other monsoon proxies, such as those invoking strong monsoon at 8 Ma (Chen et al. 2003; Zheng et al. 2004). However, a steep decrease in CIA at 8 Ma coincides with a rise in illite/ smectite (Wan et al. 2007), indicating a period of important environmental change. In addition, we plot Al/Si as a measure of the relative proportion of clays compared to quartz sand in the sediment (Fig. 11d). Clays are rich in Al and this proxy has the advantage of not being affected by the large amount of biogenic carbonate in the cores. Most of the core at ODP Site 1148 is very fine grained and no sandy or silty intervals were identified (Shipboard Scientific Party 2000), yet the Al/Si proxy shows strong temporal variations, most notably a steady decrease (i.e. less clay) since c. 8 Ma, after a period of high clay content in the Middle and Late Miocene (Fig. 11d). More sandy flux might indicate stronger physical weathering onshore after that time, consistent with the falling CIA values. We are able to gain further insight into the evolving clay mineralogy itself by reference to the CRAT proxy of Clift et al. (2008c). Clay mineralogy has been used in the past to reconstruct the changing intensity of the monsoon both in South Asia (Derry & France-Lanord 1996a) and in the South China Sea (Clift et al. 2002; Wehausen & Brumsack 2002; Trentesaux et al. 2006). CRAT is calculated based on colour spectral data and end member mixing calculations, and is designed to provide a measure of the mineralogical ratio chlorite/ (chlorite þ hematite þ goethite). The method is remote, but has been calibrated using laboratory standards under the assumption that visible range colour spectra are principally controlled by chlorite, hematite, and goethite, and clearly reflect long-term variations in core chemistry and mineralogy that cannot be linked to diagenesis. The alteration minerals hematite and goethite are largely produced by chemical weathering, whereas chlorite is indicative of physical erosion (Chamley 1989). As a result CRAT measures the relative intensities of chemical

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weathering and indicates a period of strong chemical weathering in southern China during the Middle –Late Miocene, followed by less intense weathering (and presumably a weaker summer monsoon) in the Late Miocene– Pliocene (Fig. 11a). The period of reducing chemical weathering shown by climbing CRAT values after 12 Ma correlates with increased illite/smectite at ODP Site 1146 (Wan et al. 2007) (Fig. 11b) as might be predicted. In contrast, strong chemical weathering during the Middle Miocene is shown by low illite/smectite ratios and low CRAT values. However, the high CRAT values around 16 Ma show no response in clay mineralogy. Unfortunately, the record at ODP Site 1146 does not extend far enough back in time to see whether illite/smectite was higher before 23 Ma, as might be predicted. Comparison of both CRAT values at ODP Site 1148 and illite/smectite ratios at ODP Site 1146 are hard to reconcile in a simple way with detailed CIA values. However, all three proxies suggest generally weaker summer monsoons and less weathering after the Middle Miocene. CIA does not decrease in a uniform fashion, but shows at least two positive excursions centred at c. 16– 17 and 3– 5 Ma, interpreted as periods of stronger chemical weathering and stronger summer monsoon. The period at 3–5 Ma is noteworthy in having anomalously low illite/smectite ratios, consistent with stronger chemical weathering (Fig. 11b), but this is at odds with the low CRAT value. CIA increases after 5 Ma at the same time that CRAT begins to increase, although since 3 Ma CIA has shown modest decrease at a time that CRAT values decreased, and when illite/smectite ratios were very variable. Differences between these proxies are hard to interpret. Some of the issues may relate to the fact that CRAT examines only the clay mineral fraction of the sediment, whereas CIA looks at the bulk sediment. Furthermore, CRAT values are controlled in part by goethite and hematite and do not factor in smectite, which requires higher degrees of chemical weathering to become abundant. Finally, we consider the Ti/Ca ratio, which is a good proxy for evaluating the relative influence of clastic v. carbonate sediment influx (Fig. 11e). From c. 23 to 6 Ma, the Ti/Ca ratio remained mostly at low levels, but with two short periods of higher values observed at c. 15– 17 and 10.5 – 11.5 Ma. What is most striking is the steady increase in Ti/Ca values since 4 Ma, suggestive of a rapidly increasing clastic flux to the drill site. This is consistent with shipboard core description and CaCO3 measurements (Shipboard Scientific Party 2000). This trend parallels reconstructed trends in monsoon-related foraminifera at ODP Site 1146 on the northern margin of South China Sea (Wang

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et al. 2003), as well as winter monsoon dust records from the North Pacific (Rea 1994) and from the Chinese Loess Plateau (An et al. 2001). However, enhanced post 4 Ma clastic flux to the ocean is not only a pan-Asian phenomenon that has been linked to intensifying summer monsoon (Me´tivier et al. 1999; Clift 2006), but is recognized worldwide and has been linked to fast continental erosion under the influence of variable glacial –interglacial cycles (Zhang et al. 2001).

Discussion SE Asia is a classic natural laboratory for examining possible interactions between solid Earth tectonics and climatic evolution. Many researchers have used information from sediment records in the South China Sea, as well as from bedrock onshore, in order to quantitatively model this relationship by reconstructing the timing of East Asian monsoon intensification. However, clear linkages between monsoon intensity, topographic growth and its effects on continental erosion have been hard to find because of uncertainties in all three processes. Results from our study suggest that combinations of geochemical and mineralogical data derived from ODP Site 1148, with sediment budgets from the Pearl and Red Rivers show some correlation, especially in the Pleistocene and Early Miocene, suggesting that climate variability is a controlling factor on continental weathering in SE Asia. The role of topographic uplift is harder to constrain because of the uncertainties in Tibetan and SE Asian topographic growth (Clark et al. 2005; Harris 2006; Schoenbohm et al. 2006a) and the possible influences of drainage reorganization in governing the flux of sediment to the Gulf of Tonkin. Reconstructing the stratigraphic evolution of the SH-Y and Qi Basins is important to studies of monsoon-tectonic coupling relationships for several reasons. These two basins together form one of the largest sedimentary masses in SE Asia and record the erosion flux from a major drainage that cuts the flank of the Tibetan Plateau over a long period of time. Uplift of SE Tibet and Yunnan might therefore be expected to have driven faster mass fluxes as the gorges of the upper Red River were cut. Although this is a region of strong summer monsoon rains the heaviest rains are closer to the coast, not in the areas of strongest surface deformation. A stronger monsoon might be expected to cause stronger run-off, higher erosion and faster sedimentation rates. This area contrasts with the erosional links proposed for the frontal ranges of the Himalaya where the steep topographic gradient results in a close coupling of

climate, exhumation and structure (Hodges et al. 2004; Thiede et al. 2004; Clift et al. 2008c). In contrast, the topographic gradient of the edge of the Tibetan Plateau is much more gradual along its southeastern flank.

Tectonic and stratigraphic evolution Although the nature of the oldest sedimentary formations within the SH-Y and Qi Basins is still uncertain, there is little doubt that deposition had started during the Eocene–Oligocene (Zhong et al. 2004; Clift & Sun 2006). The basin began to subside strongly after c. 34 –35 Ma when motion on the RRFZ first started (Gilley et al. 2003). This motion and its associated transtensional faulting together controlled formation of the SH-Y Basin as a pull-apart. Motion on the RRFZ probably caused the pre-rift formations to be uplifted, deformed and eroded, at least locally (Figs 3 & 7). This effect is best observed in the northern part of the SH-Y basin. Relatively stationary depocentres during the Oligocene–Early Miocene suggest that most of sediments delivered from the Red River were trapped in the northern and central SH-Y Basin, while little sediment was reaching the Qi Basin. Active faulting and basin subsidence continued until c. 21 Ma and was followed by slower thermal subsidence after an inversion event focused in the northern SH-Y Basin, younging to c. 15 Ma in the south. The cessation of seafloor spreading and the end of left-lateral motion on the RRFZ together resulted in an inversion event in the early Middle Miocene. These processes correlate well with the change in the basin from transtensional to transpressional character. As a result the whole basin was strongly inverted, deformed and eroded before 15 Ma. A significant portion of the pre-uplift formations in the northwestern SH-Y Basin was removed following strong uplift and erosion (Figs 3 & 7). For this reason, very little of the sediment delivered by the Red River was deposited in that part of the system at that time. Instead, sediments were bypassed to the centre and southeastern end of the basin where the basin floor was still deep and where accommodation space allowed preservation. However, the Qi Basin did not experience Mid-Miocene inversion, presumably because of its distance from the RRFZ. After inversion, the SH-Y Basin gradually subsided again. Low-angle shelf edge trajectories observed in some places demonstrate that less accommodation space was created after 15 Ma, consistent with reconstructions of basement tectonic subsidence (Clift & Sun 2006). None the less, the integrated basin-wide sedimentary budget shows faster sedimentation at 11–15.5 Ma compared to the Early Miocene.

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Influence of Hainan Clinoforms observed in Figures 4 and 5 show that the northwestern SH-Y Basin was nearly filled after c. 5 Ma, so that most of the younger sediments must have overspilled not only to the SE, but partly into the Qi Basin, as they did during the Last Glacial Maximum (c. 20 ka). Sediments derived from the Red River and eroded from Hainan island were deposited together in the northwestern half of the Qi Basin as a large prograding clastic wedge. The high-angle shelf edge trajectory and the direction of progradation observed in southern Hainan (Fig. 6) suggest that significant volumes of sediments eroded from the island were delivered to the Qi Basin. The increase in sediment supply from that area may be a response to the tectonically driven uplift of the Hainan island, linked to magmatism during the Pleistocene (Tu et al. 1991; Flower et al. 1998). In addition, faster erosion in Hainan and within the Red River Basin driven by an intensified summer monsoon may be responsible for some of the increased erosion. Finally we consider that progressive surface uplift in SW China (Yunnan) and northern Vietnam within the headwaters of the Red River during the Pliocene (Schoenbohm et al. 2006a) may have driven faster erosion by causing enhanced incision of the Red River.

Chemical weathering and climate change Intensification of the East Asian monsoon may be one of the most important factors in controlling continental weathering processes. On the continents precipitation, temperature and vegetation are the primary controls on both chemical weathering and physical erosion (White & Blum 1995; Edmond & Huh 1997; West et al. 2005). Over geological timescales, a measure of this is preserved in the chemistry and mineralogy of sediments transported by rivers. Water-mobile elements are easily removed from weathering products and whereas more stable elements tend to be relatively enriched. Based on these principles, we now combine information derived from the sedimentary budget and from geochemical analysis to assess how changing monsoon strength may have influenced erosion within the Red River drainage since the Oligocene. Figure 10c shows that sedimentation flux from the Red River Basin initially increased after c. 29 Ma and subsequently fell again after 21 Ma. Unfortunately, there is no weathering record predating 24 Ma. Clay mineralogical records from ODP Site 1148 show monsoon strengthening after c. 22 Ma (Clift et al. 2008c). If this is true then the higher erosion rates at 21–28 Ma largely predate a strong monsoon, while the first period of strong summer monsoon (22– 17 Ma) correlates with a

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period of reduced sediment flux to the SH-Y Basin (Fig. 10a). The increase in sediment flux into the basin at 21–29 Ma was thus more likely to have been triggered by the onset of topographic uplift and exhumation related to the Red River Shear Zone, which started at c. 34– 35 Ma (Leloup et al. 2001; Gilley et al. 2003). A decrease in sedimentation rate after c. 21 Ma may indicate erosion of the early topography. Minimum CRAT values and low illite/smectite ratios from c. 15 –10 Ma, together with steady CIA values (Fig. 11a, b, c) indicates that chemical weathering was strong in a climate of intensified summer monsoon rain (Wan et al. 2007). This time correlates with a period of moderately increased sedimentation rates (Fig. 11f ). A positive link between erosion rates and monsoon intensity is suggested at that time, not least because the RRFZ became inactive after 15 Ma, so that the increased rates of erosion are the opposite of those predicted if tectonic forces were the dominant erosion control. Decreased CIA values after 8 Ma indicates that chemical weathering weakened after this time. Similarly, rising CRAT values, high illite/smectite and decreased sedimentation rates all point to weaker chemical weathering in a drying monsoon climate after 8 Ma, correlating with a fall in clastic flux. This drop in sedimentation rates during the Late Miocene parallels similar synchronous trends recognized on the Bengal Fan (Burbank et al. 1993). However, unlike that study we suggest a positive correlation between erosion and monsoon strength because 8 Ma now appears to mark a time of summer monsoon weakening not intensification as previously believed (Derry & France-Lanord 1996b).

Tibetan gorge incision Although the RRFZ was reactivated after 5 Ma, albeit in a reverse, dextral sense (Schoenbohm et al. 2006b), the degree of active shear in the Red River drainage never regained the rates seen in the Middle –Early Miocene. However, progressive uplift of eastern Tibet, driving gorge incision along the edge of the plateau may have been an influence on sediment flux to the SH-Y and Qi basins. Thermochronological work in Yunnan and Sichuan in SW China indicates accelerated surface uplift there starting around 11 Ma (Clark et al. 2005). However, this was a time of slower sedimentation in the SH-Y and Qi basins, indicating either that the drier monsoon was the greater influence on erosion or that the gorge incision in SW China was not feeding sediment to the Red River. Reconstructions of the rivers around the eastern Himalayan syntaxis suggest that capture of the Yunnan rivers away from the Red River was completed

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before that time (Clark et al. 2004; Clift et al. 2006a), so that any sediment pulse would have been diverted into the East China Sea, consistent with our sediment budgets. Surface uplift in northern Vietnam postdates that in Yunnan, being mostly Pliocene in age and reflecting gradual growth of topography to the SE (Schoenbohm et al. 2006a). This phase of tectonism and associated gorge incision correlates well with the pulse of sediment seen in the SH-Y and Qi basins. The Ti/Ca curve from ODP Site 1148 shows an increase at this time despite the fact that the Pearl River Basin is less affected by topographic uplift. Intensification of the summer monsoon since c. 4 Ma in the South China Sea region may be the dominant control on increasing erosion within this time period (Wan et al. 2006), whereas rock uplift now controls the patterns of erosion within the Red River Basin itself (Clift et al. 2006b). Our climate reconstruction is consistent with regional compilations of increased sediment flux in the Middle Miocene (Clift 2006), although our records favour an earlier start to higher sediment flux, during the Early Miocene. The higher rates of sediment flux at 29–21 Ma are unique to the Red River system and support a local tectonic rather than regional climatic trigger. Our apparent initial monsoon intensification after c. 22 Ma is much earlier than the commonly cited 8 Ma monsoon intensification (Kroon et al. 1991; Prell et al. 1992; Zheng et al. 2004), but is consistent with the revised summer monsoon model of Clift et al. (2008c). Our estimate is older than the c. 15 Ma intensification suggested by Wan et al. (2007), but is consistent with palynology and facies information from China (Sun & Wang 2005). Decreasing humidity from c. 8 –4 Ma and especially a rapid drop between 5 and 4 Ma, charted by falling CIA values (Wei et al. 2006), falling kaolinite contents (Wan et al. 2007) and rising CRAT values (Clift et al. 2008c) (Fig. 10) testify to a weakening monsoon. This resulted in less physical erosion in the mountainous Red River basin and reduced chemical weathering in the flatter Pearl River Basin. Chemical weathering is further reduced by falling global temperatures since the Middle Miocene (Zachos et al. 2001). This change is shown by low sedimentation rate (Fig. 11f) and by decrease in the Al/Si ratios (Fig. 11d). The period of c. 3–4 Ma is marked by falling CRAT values, higher illite/smectite ratios and lower CIA, which are not all in accord regarding the long-term change in summer monsoon strength. Summer monsoon strength varies rapidly over millennial timescales at this time and a longer duration pattern is hard to discern (Clift & Plumb 2008). This period is also accompanied by decreases

in the Al/Si and CIA ratios whereas the Ti/Ca increased, suggesting enhanced coarser influx, as well as stronger chemical weathering. Rapid increase in the Ti/Ca ratio after c. 4 Ma, especially after c. 2.7 Ma indicates an increase in clastic sediment influx relative to carbonate sediments, which was probably caused by enhanced continental physical weathering driven by the transition between glacial-interglacial climate states (Zhang et al. 2001).

Conclusions Tectonically driven surface uplift of eastern Tibet has commonly been linked to enhancement of the East Asia monsoon. Both these processes have the potential to increase the rates of continental erosion, which should be reflected in the volumes and composition of sediments in river deltas. This study shows that a combination of sedimentary budgets derived from regional seismic stratigraphic and geochemical data can be employed to compare weathering regimes and erosion rates over tectonic time periods .20 Ma. A combination of proxies allows us partially to reconstruct the history of monsoon climate change since 24 Ma. Our work confirms that the SH-Y and Qi basins formed after c. 50 Ma and especially subsided rapidly after c. 34 Ma, coincident with the onset of motion on the RRFZ. The SH-Y Basin experienced two inversion phases that occurred at c. 34 and c. 15.5 Ma, while the Qi Basin seemed not to be affected by these events. 34 –17 Ma motion on the RRFZ correlates with a period of faster sedimentation in the SH-Y Basin. This is despite the initial intensification of the monsoon dating from only c. 22 Ma. We conclude that tectonic forces are dominant in controlling erosion at that time. Geochemical data suggest that chemical weathering has generally decreased since c. 25 Ma, while physical erosion became stronger. A shift to more physically eroded chlorite and increasing sedimentation rates after c. 15.5 Ma points to stronger rains and stronger physical erosion between 15.5 and 10 Ma. In this period climate appears to dominate as the primary erosional control, as motion on the RRFZ had ceased. The period from 10 to 4 Ma saw a reduction in chemical weathering and sediment flux, correlating with a time of weakening summer monsoon. However, the transition to glacialinterglacial climates, surface uplift in northern Vietnam and Hainan island and stronger summer monsoons, at least during the interglacial periods, since 4 Ma correlates with a switch back to stronger erosion of the source rocks, especially physical erosion, which in turn raised the clastic influx into the basins.

TECTONIC AND MONSOON EVOLUTION IN SE ASIA We thank the Natural Environment Research Council (NERC) in the United Kingdom and the College of Physical Sciences at the University of Aberdeen for funding and support for this project. We particularly thank BP Exploration for release of new seismic data to our project. We thank PetroVietnam, the Chinese National Offshore Oil Company (CNOOC), and Integrated Ocean Drilling Program (IODP) for additional supporting data. Seismic Micro-Technology Inc. provided use of the KingdomTM seismic interpretation software. We also wish to thank David Heslop, Alan Roberts, Nick Kusznir, Ro¨hl Ursula, Vera Lukies, Prof. Mai Thanh Tan and Tran Thi Kieu Hoa for technical advice.

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Geochemical records in the South China Sea: implications for East Asian summer monsoon evolution over the last 20 Ma SHIMING WAN1,2*, PETER D. CLIFT3,4, ANCHUN LI1, TIEGANG LI1 & XUEBO YIN1 1

Key Laboratory of Marine Geology and Environment, Institute of Oceanology, Chinese Academy of Sciences, Qingdao 266071, China

2

Institute of Geology and Geophysics, Chinese Academy of Sciences, Beijing, 100029, China 3

School of Geosciences, University of Aberdeen, Meston Building, Kings College, Aberdeen, AB24 3UE, UK 4

South China Sea Institute of Oceanology, Chinese Academy of Sciences, 164 Xingang Road, Guangzhou, 510301, China *Corresponding author (e-mail: [email protected])

Abstract: We reconstruct past changes in the East Asian summer monsoon over the last 20 Ma using samples from Ocean Drilling Program (ODP) Site 1146 of Leg 184 in the northern South China Sea based on the major (Al, Ca, Na, K, Ti, etc.) and trace element (Rb, Sr, and Ba) geochemistry of terrigenous sediments. This study and combined review suggests that the long-term evolution of the East Asian summer monsoon is similar to that of the Indian summer monsoon, but distinct from the East Asian winter monsoon. Generally, the Asian summer monsoon intensity has decreased gradually from its maximum in the Early Miocene. In contrast, the Asian winter monsoon shows a phased enhancement since 20 Ma bp. Moreover, our study shows that the long-term intensities of the Asian summer and winter monsoons may have different forcing factors. Specifically, the winter monsoon is strongly linked to phased uplift of Tibetan plateau and to Northern Hemispheric Glaciation. In contrast, global cooling since 20 Ma bp may have largely reduced the amount of water vapour held in the atmosphere and thus weakened the Asian summer monsoon.

The Asian monsoon system, comprising the East Asian and Indian (or South Asian) subsystems, exerts a dominant influence on Asian climate, and its evolution plays a significant role in our understanding of the regional and global climate (Webster et al. 1998; B.Wang et al. 2003). Progressive surface uplift of the Tibetan Plateau, mostly following the India-Asia collision, has been inferred as a primary cause of monsoon initiation and/or intensification (e.g. Ruddiman & Kutzbach 1989; Raymo & Ruddiman 1992; An et al. 2001; Zheng et al. 2004; Harris 2006; Wan et al. 2007), although other mechanisms have also been invoked. The link between monsoon strength and Tibetan elevation is intriguing, but has not yet been demonstrated, not least because the growth of the plateau remains enigmatic and long-term records of monsoon evolution are rare. One way of understanding the proposed climate –tectonic interactions is to define the history and timing of major tectonic and climatic changes in Asia, especially the development of the Asian monsoon. Until now, most palaeoclimatic studies which permit the long-term reconstruction of the Asian palaeo-monsoon since the Late Miocene have

focused on records in the loess–palaeosol sequences of the Chinese Loess Plateau (e.g. Ding et al. 1999; An et al. 2001; Qiang et al. 2001; Guo et al. 2002; Jiang & Ding 2008), palaeobotanical records in the Chinese mainland (Sun & Wang 2005), sediment records in the South China Sea (Clift et al. 2002, 2008b; Chen et al. 2003; Jia et al. 2003; Jian et al. 2003; P. Wang et al. 2003; Tian et al. 2004; Zheng et al. 2004; Hess & Kuhnt 2005; Clift 2006; Wan et al. 2006, 2007; Wei et al. 2006), and sediment records in the Arabian Sea and Bengal Fan (Kroon et al. 1991; Prell et al. 1992; Burbank et al. 1993; Derry & France-Lanord 1996; Gupta et al. 2004; Huang et al. 2007; Clift et al. 2008b), as well as palaeosol sequences in the Himalayan foreland basin and Arabian Peninsula (Quade et al. 1989, 1995; Dettman et al. 2001; Sanyal et al. 2004; Behrensmeyer et al. 2007). In comparison with the East Asian winter monsoon and Indian summer monsoon, the long-term variation of the East Asian summer monsoon from the Early Miocene to the present is not well known. A low-resolution record of long-term chemical weathering in the South China Sea drainage area has shed some light on the evolution of East Asian

From: Clift, P. D., Tada, R. & Zheng, H. (eds) Monsoon Evolution and Tectonics –Climate Linkage in Asia. Geological Society, London, Special Publications, 342, 245–263. DOI: 10.1144/SP342.14 0305-8719/10/$15.00 # The Geological Society of London 2010.

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summer monsoon (Wei et al. 2006), as has recent high-resolution studies based on colour spectral analysis (Clift et al. 2008b). However, summer monsoon intensity during some crucial times of proposed change, such as 17 –15 Ma, around 8 Ma, and 3–4 Ma remains obscure. Furthermore, previous studies have been largely concerned with regional monsoon variability, rather than the entire Asian monsoon system. In order to understand monsoon evolution and its driving forces better, it is important to determine whether the summer and winter monsoons evolve in phase or not, whether the East Asian and Indian monsoons are coupled over tectonic timescales, and whether the intensity of the entire Asian monsoon system correlates with the phased uplift of the Tibetan Plateau or is controlled by other forcing factors. This study aims to address these questions through a much higher resolution long-term study than those already recently performed in the Loess Plateau (Jiang & Ding 2008) or in the South China Sea (Wei et al. 2006; Clift et al. 2008b). We reconstruct Asian summer monsoon intensity through examination of the degree of silicate chemical weathering. Silicate weathering strongly affects the major-element geochemistry of siliciclastic sediments (e.g. Nesbitt & Young 1982; McLennan 1993), for example larger cations (Al, Ba, Rb) remain fixed in the weathering profile whereas smaller cations (Ca, Na, Sr) are selectively leached (Nesbitt et al. 1980). These chemical signatures are ultimately transferred to the sedimentary record, thus providing a useful tool for monitoring source area weathering conditions. The Nd isotopic compositions of sediments from ODP Site 1148 indicate that the source of sediment

since 23 Ma is generally stable and mainly from South China (Clift et al. 2002; Li et al. 2003). Rainfall in southern China is dominated by the East Asian summer monsoon (Ding 1994), which controls the humidity in this region. Variations in the degree of chemical weathering in southern China should mostly reflect changes in the East Asian summer monsoon because the intensity of chemical weathering is largely controlled by humidity (Berner & Berner 1997). In this paper, we extract a series of geochemical proxies for palaeoclimate derived from the major and trace element (Rb, Sr, and Ba) composition of siliciclastic sediments recovered from the South China Sea in order to: (1) reconstruct the long-term history of chemical weathering intensity in South China and use this to infer changes in the intensity of the East Asian summer monsoon over the last 20 Ma; (2) compare the similarities and differences of the long-term history of the East Asian summer and winter monsoons, as well as the Indian summer monsoon since the Early Miocene; and (3) discuss the possible links between the Asian monsoon system and the uplift of the Tibetan Plateau as well as other controls such as global ice volume.

Materials and methods ODP Site 1146 is located at 19827.400 N, 116816.370 E, at a water depth of 2092 m, within a small rift basin on the mid-continental slope of the northern South China Sea (Fig. 1). Three holes were cored to a sub-seafloor depth of 643 m composite depth (mcd) (Wang et al. 2000). For this study,

Fig. 1. Location map showing the geographic features of Asia and the surrounding oceans (modified after Zheng et al. 2004).

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a total of 274 samples were sampled continuously at 3 m intervals from 02327.50 mcd and 499.84– 642.44 mcd and at 1.5 m intervals from 327.50– 498.94 mcd. The lithology of the recovered section changes greatly and is dominated by hemipelagic fine-grained terrigenous materials and nannofossil carbonate ooze. The middle Miocene through Pleistocene section is characterized by relatively carbonate-rich, hemipelagic nannofossil clay. In contrast, the Pleistocene sediments are composed of greenish-grey nannofossil clay that is relatively enriched in quartz, feldspar and chlorite (Wang et al. 2000; Wan et al. 2007). Terrigenous materials primarily comprise quartz, feldspar, and clay minerals and account for up to 99% of the clastic fraction. Carbonate contents range from 13% to 64% throughout the section (Wan et al. 2007). The chronostratigraphic framework for ODP Site 1146 was established on the basis of the magnetostratigraphy and biostratigraphy (Wang et al. 2000) and then interpolated linearly between control points. ODP Site 1146 is characterized by moderate accumulation rates (average c. 3.3 cm ka21), but these increase sharply (9.5 cm ka21) during the Pleistocene. This sediment sequence spans c. 19.5 Ma. We sampled with a resolution of about 70.9 ka. Geochemical analysis of major elements, as well as Rb, Sr, and Ba concentrations was performed on bulk organic-, and carbonate-free sediments. Organic matter, calcite, and even Fe –Mn oxides were removed by treating with 10% H2O2 at 60 8C for 1 hr and 0.5 N HCl at 60 8C for 2 hrs, respectively, in order to isolate the granular siliciclastic particles. Finally, the sediments were rinsed with deionized water three times and dried at 80 8C, before being ground into powder for elemental composition analysis. This chemical pre-treatment effectively removes most of the biogenic/authigenic materials in the sediments, including carbonate, Fe –Mn oxides and organic materials, except for biogenic opal (Li et al. 2003; Wei et al. 2006). According to the initial reports of ODP Leg 184 (Wang et al. 2000), there are very low contents of authigenic Fe oxide and pyrite (average 1%) in the sediments of ODP Site 1146, with little dolomite, glauconite, zeolite and barite. Moreover, the biogenic opal contents in the sediments are very low (average 3%) (Wan 2006) and many of the elements, such as Al, K, Na, Ca, Mg, Ti and rare earth elements are undetectable in biogenic opal (Wei et al. 2006). A similar chemical procedure was applied in earlier studies (Li et al. 2003; Wei et al. 2004, 2006) and proved that most of the elements that are not concentrated in authigenic components show similar variation patterns both in treated and bulk sediments and thus the ratios between these elements are identical to the

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corresponding ratios of the detrital components (Wei et al. 2004, 2006). In addition, four samples (at 181.80 mcd, 384.35 mcd, 411.37 mcd, 513.69 mcd) from ODP Site 1146 were randomly selected and separated into five distinct grain size fractions (,2 mm, 2 –4 mm, 4–8 mm, 8 –16 mm, .16 mm) by gravity setting in deionized water. Then the pre-treated sediments were digested by concentrated HF þ HNO3 þ HClO4 mixture, and the major oxide (except SiO2) and trace element concentrations were determined using a Thermo iCAP6300 ICP-AES and a Perkin-Elmer ELAN DRC II ICP-MS, respectively, at the Institute of Oceanology, Chinese Academy of Sciences, Qingdao. Detailed analytical methods are described in Li et al. (2003). Several USGS and Chinese rock and sediment standards (AGV-2, BCR-2, BHVO-2, GBW07316, GBW07315, and GBW07333) and blanks were repeatedly digested and measured along with the samples to monitor the quality of ICP-MS and ICP-AES measurements, and the results were generally within the range of +10% of the certified values. The analytical precision is generally better than 1–2% for major elements and 1 –3% for trace elements Rb, Sr, and Ba.

Results and discussions Elemental compositions and hosted mineral phases The aluminosilicate fractions of sediments from ODP Site 1146 consist mainly of SiO2 (not measured), Al2O3, K2O, Fe2O3, and MgO, with low concentrations of Na2O, CaO, TiO2, P2O5, and MnO (Fig. 2). The major and trace (Rb, Sr, Ba) element compositions of sediments from ODP Site 1146, as well as the Pearl, Red, Mekong, Yangtze and Yellow Rivers, offshore of SW Taiwan, West Philippine Sea (WPS) and loess were all normalized to upper continental crust (UCC); (Taylor & McLennan 1985) and are shown in Figure 3. Most of the elements, including Ca, K, Mg, Mn, Na, P, Sr, and Ba, are more or less depleted in the sediments from ODP Site 1146, whereas Ti and Rb are relatively enriched. In contrast, Al and Fe in sediments of different ages oscillate between depletion and enrichment. The obviously lower UCC-normalized concentrations of Mg, Mn, Na, and P in sediments from ODP Site 1146 compared to those from the Yangtze and Yellow Rivers, loess, the West Philippine Sea, offshore SW Taiwan, and even the modern Pearl River suggests stronger chemical weathering in the source-areas of ODP Site 1146 at the time of deposition compared to the other regions at present.

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Fig. 2. Variation of mineralogy (Wan et al. 2007), major elements and Rb, Sr, and Ba abundances, mass accumulation rates (MAR), mean grain size of siliciclastic sediments at ODP Site 1146 in the northern South China Sea since about 20 Ma bp. For comparison, the global eustacy curve (Haq et al. 1987) and Nd isotopic variations in sediments at ODP Site 1148 (Li et al. 2003) are also plotted.

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Fig. 3. Spider plot showing comparisons of elemental concentrations in the siliciclastic sediments at ODP Site 1146 in the northern South China Sea with those from ODP Site 1148 (Wei et al. 2006), Pearl, Red, Mekong (Liu et al. 2007), Yangtze, Yellow River (Yang et al. 2004), SW Taiwan (Selvaraj & Chen 2006), West Philippine Sea (WPS) (Shi & Chen 1995), and Loess (Yang et al. 2006a). Because of the large amount of data, the full compositional range of sediments at ODP Site 1146 is shown by shaded area.

Variations in elemental concentrations over the last 20 Ma are shown in Figure 2. The long-term changes of Al2O3, K2O, Rb, and Ba show similar trends. In general, these elements remain relatively stable downhole, but there is a significant decrease in concentrations between 5 and 3 Ma. In contrast, Ca, Na, and Sr exhibit a generally increasing trend. Moreover, Mg, Fe, and Mn show slightly lower concentrations between 11 and 3 Ma. The long-term variation in TiO2 between 11 –3 Ma is similar to Ca, Na, and Sr, but with a decreasing trend for the last 3 Ma. Different from the other elements, P2O5 shows a relatively stable trend throughout the last 20 Ma. Generally, each element is concentrated in one or several specific mineral phases, and different kinds of mineral are often composed of different associated elements, such as albite, which consists of Si, Al, O, and Na, whereas anorthite consists of Si, Al, O, and K. The chemical weathering of rocks is effectively expressed in two aspects. One is the transformation of primary minerals to secondary mineral phases, and the other is the leaching, mobilization and precipitation of mobile elements

during weathering (e.g. Nesbitt et al. 1980). For example, Na, Ca and Sr are preferentially and rapidly lost compared with Al and Ti during the weathering of plagioclase to illite. On the basis of the mineral assemblages, contents, and chemistry in terrigenous-derived sediments at ODP Site 1146 (Wan et al. 2007, 2008), the proportions of major elements hosted in specific mineral phases were calculated by multiplying the concentration of the element in an individual mineral by the weight percent value of the individual components in the terrigenous sediments at ODP Site 1146 (Table 1). On average, about 55% of Al in the sediments at ODP Site 1146 resides in illite, with only 14% in smectite, 12% in kaolinite, 9% in chlorite, and 8% in plagioclase. Among the alkalis and alkaline elements, Ca, Na, and Sr mainly concentrate in plagioclase and smectite, whereas K mostly concentrates in illite, but with smaller amounts in K-feldspar. Rb and Ba may substitute for K in the lattice of silicate minerals and is generally also common in illite and K-feldspar (e.g. Nesbitt et al. 1980). In contrast, Mg, Fe, and possibly Mn are primarily enriched in chlorite and to a lesser

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Table 1. The proportions of occurrence of major elements and Rb, Sr, Ba hosted in specific mineral phases (%) Minerals

Si

Al

Ca

Fe

K

Mg

Na

Quartz K-feldspar Plagioclase lllite Smectite Kaolinite Chlorite

54 1 4 20 12 4 4

0 1 8 55 14 12 9

0 0 62 0 38 0 0

0 0 0 0 18 0 82

0 6 0 94 0 0 0

0 0 0 0 22 0 78

0 0 63 0 37 0 0

Mn

Rb

Sr

P

Ba P

P P P

P P

P

Note: The symbol ‘P’ is the abbreviation of ‘possible’, indicating that the designated mineral possibly contains this element. Minerals components used in the calculations are quartz (SiO2); K-feldspar (KAISi3O8); plagioclase (CaAI2Si2O8); illite (K0.75AI2.75Si3.25O10(OH)2); smectite (Na0.33/2Ca0.33/2AI1.4Mg0.5Fe0.1Si4O10(OH)2-H2O); kaolinite (AI2Si2O5(OH)4); chlorite (Mg4Fe1AI2Si3O10(OH)4). The average contents of minerals in terrigenous sediments of ODP Site 1146 are as follows quartz 22%, K-feldspar 1.3%, plagioclase 6.7%, illite 35%, smectite 17.5%, kaolinite 7%, chlorite 10.5%.

extent in smectite (Table 1). These different associations of elements are consistent with the above result deduced from long-term variations of elements.

Elemental mobility and weathering trends Relative mobility of elements refers to changes in concentration of elements in the alteration profile relative to those in the parent rocks (Nesbitt 1979). Here we take the UCC as a reference for the parent rocks. In order to define relative mobility, an immobile element is needed. Ti is often considered a conservative element and is chosen as an immobile constituent. Thus the relative mobility of any given element X is given by [(X/Ti)sample/ (X/Ti)UCC21]  100 (Nesbitt 1979). The degree of chemical weathering is generally estimated by the chemical index of alteration (CIA) (Nesbitt & Young 1982). CIA is defined as 100  Al2O3/ (Al2O3 þ CaO* þ Na2O þ K2O) (molecular proportions, with CaO* being the CaO content in the silicate fraction of the sample). As a result the behaviour of an element during chemical weathering can be assessed by plots of CIA against the relative mobility of that element. As shown in Figure 4, with increasing chemical weathering (increasing CIA), the depletion of Ca, Na and Sr becomes more prominent, suggesting that these three elements are highly mobile and are preferentially removed from rocks during weathering processes. The strong mobility of Ca, Na and Sr is caused by heavy dissolution of plagioclase from source rocks under strong weathering. Other elements, including K, Mg, Fe, Mn, P, Rb, and Ba appear relatively immobile (Fig. 4). Their variation indicates they were generally less affected by weathering processes, possibly because of their rapid precipitation in secondary clay minerals that prevented loss during weathering (Nesbitt & Markovics 1997).

The weaker mobility of K, Rb, and Ba relative to Ca, Na and Sr indicates that the dissolution of K-feldspar is insignificant in comparison with plagioclase. In contrast, Al is progressively enriched as the CIA values increase (Fig. 4). Apparently, more highly aluminous fine clays from the soil zones were translocated and thus more Al was introduced to these samples as the degree of chemical weathering increased (Nesbitt & Markovics 1997). Weathering trends in sediments from ODP Site 1146 can be clearly observed on Al2O3 – (CaO* þ Na2O)– K2O (A –CN –K) ternary diagrams (e.g. Nesbitt & Young 1989). As shown in Figure 5a, all samples from ODP Site 1146 are arranged as a group and subparallel to the A-CN line, suggesting a similar composition to the parent rocks since 20 Ma bp. Moreover, sediments from ODP Sites 1146 and 1148 almost overlap in the A-CN-K diagram (Fig. 5b), implying that the two drilling sites have had a similar provenance since the Early Miocene. In addition, the trend of silicate weathering displays the preferential leaching of CaO and Na2O and enrichment of Al2O3, while K2O contents remain constant. Therefore, we infer that plagioclase was selectively weathered first, while K-feldspar remained intact, even under a range of different weathering conditions, at least since the Early Miocene. Nesbitt & Markovics (1997) also observed that K leaching from the parent rock during the early and intermediate weathering stages was not strong. In addition, the chemical compositions of sediments from the Yellow River, Loess Plateau, the Yangtze River, offshore SW Taiwan, and the Red, Mekong, and Pearl Rivers also fall on the same weathering-trend as those from ODP Site 1146 (Fig. 5a), revealing a gradual increase in weathering intensity from the Yellow to the Pearl River basin. The original parent rock composition can be predicted by the point at which the weathering-trend

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Fig. 4. Diagrams showing how the relative mobility (RM) of a series of different elements changes with increasing chemical alteration index (CIA). With more intense chemical weathering (increasing CIA), Ca, Na and Sr becomes more depleted, whereas Al and Rb are progressively enriched. In contrast, the other analysed elements (Fe, Mg, K, Mn, P and Ba) show no coherent change with chemical weathering.

line intersects the plagioclase-K-feldspar junction (Nesbitt & Young 1989). It is evident that the parent rocks of major rivers in East and South China, the loess source regions, and Taiwan are similar in their major element composition and lie very close to that of UCC. In contrast, the parent rocks of sediments in the West Philippine Sea are more like an andesite (Fig. 5a).

Proxy implications for chemical weathering in South China As discussed above, the behaviour of different elements is very different during chemical weathering. Weathering products comprise the majority of the detrital components in the marine sediments. As a result, the chemical signatures of each source region are ultimately transferred to the sedimentary records (e.g. Nesbitt & Young 1982). Consequently, we can select a series of elemental ratios and indices to reconstruct variations in the intensity of chemical weathering, and thus to constrain the development of the East Asian monsoon from the detrital sediments. The chemical index of alteration (CIA) value has been long believed to give a quantitative measure for chemical weathering intensity (Nesbitt & Young 1982). We define the term ‘chemical

weathering intensity’ used in this study as ‘the degree of chemical alteration of the host rocks in the drainage area due to chemical weathering’. Monitoring plagioclase weathering alone also yields additional information on silicate weathering and the plagioclase index of alteration (PIA) can be calculated as 100  (Al2O3 – K2O)/(Al2O3 þ CaO* þ Na2O–K2O) in molar proportions (Fedo et al. 1995). A higher PIA value indicates a higher degree of plagioclase weathering. As plagioclase is preferentially more rapidly weathered compared to K-feldspar during silicate weathering (Fig. 5), higher K2O/(CaO* þ Na2O) molar ratios should imply more advanced chemical weathering (Liu et al. 2007). In addition, increased chemical weathering intensity rapidly leaches Sr compared to Rb and Ba (Fig. 4; Nesbitt & Young 1982), so that the Rb/Sr and Ba/Sr ratio increases with higher weathering degree. Aluminium tends to be enriched in weathering products, whereas titanium usually behaves as a conservative element in a weathering profile. Therefore, enhanced chemical weathering would be expected to result in higher Al/Ti ratios. The Al/Ti ratio of sediments at ODP Site 1144 and 1148 has been previously used as a proxy of chemical weathering in South China by Wei et al. (2004, 2006). These selected element ratios also display positive correlation with CIA (Fig. 6).

252 S. WAN ET AL. Fig. 5. Al2O3 –(CaO* þ Na2O)– K2O (A– CN–K) ternary diagram of siliciclastic sediments from ODP Site 1146. Data of detrital sediments from ODP Site 1148 (Wei et al. 2006), Pearl, Red, Mekong (Liu et al. 2007), Yangtze, Yellow River (Yang et al. 2004), SW Taiwan (Selvaraj & Chen 2006), Loess (Yang et al. 2006a), West Philippine Sea (WPS) (Shi & Chen 1995) were plotted for comparison. Upper continental crust (UCC) (Taylor & McLennan 1985), North American shale composite (NASC) (Gromet et al. 1984), average composition of granodiorite (Grano.), granite, andesite, and basalt in Eastern China (Yan & Chi 1997) also plotted as a reference. Data of ODP Site 1148 (Wei et al. 2006) and five grain size fractions of four samples from ODP Site 1146 are shown in parts (a), (b) and (c), respectively. Arrows indicate predicted weathering trend exhibited by the sediments from South China Sea. The diagram also represents the fields of idealized minerals: Pl, plagioclase; Ks, K-feldspar; Sm, smectite; Mu, muscovite; ILL, illite; Ka, kaolinite; Chl, chlorite; Fel, feldspar; Gi, gibbsite.

ASIAN SUMMER MONSOON EVOLUTION Fig. 6. Correlation of mean grain size, PIA, Rb/Sr, Ba/Sr, Al/Ti, K2O/(CaO þ Na2O) with CIA of siliciclastic sediments at ODP Site 1146. The circle symbol corresponds to the left y-axis and the cross symbol corresponds to the right y-axis.

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In addition to chemical weathering in the source regions, other factors may also influence the elemental ratios in detrital sediments, such as provenance changes, hydraulic sorting, as well as diagenesis and/or metasomatism after burial (Fralick & Kronberg 1997). To a large extent, the elemental composition of the sediments is dominated by the constituent mineral phases and especially the proportion of quartz and feldspar relative to phyllosilicate minerals, which is strongly influenced by the variation in grain size (Wan et al. 2007). It is not surprising that the mean grain size of terrigenous sediments at ODP Site 1146 shows a moderate positive correlation (R2 ¼ 0.37) with CIA (Fig. 6). In general, the grain-size distribution is controlled by weathering in source areas and sedimentary sorting under hydraulic transport to the drill site. Regional sea level in the South China Sea gradually fell from 15 Ma to 10 Ma and then gradually rose until about 5 Ma (Fig. 2; Haq et al. 1987; Li et al. 2005). As a result, sedimentation rates on the deepwater continental margins would be expected to have peaked around 10 Ma if the sea level change controls sedimentation, assuming that the sediment yield in the drainage basin remained constant. However, between 15 –3 Ma there is no remarkable change in any of the MARs in total terrigenous materials, quartz, feldspar and clay minerals at either ODP Site 1146 or 1148 (Clift 2006; Wan et al. 2007). Sediment budgets for the major basin systems in the Asian marginal seas also show that the total flux of clastic material from Asia into the surrounding marginal seas is mostly controlled by the erosional response to monsoonal precipitation and uplift of Tibet, not to global eustacy (Clift 2006). In order to constrain the effects of hydraulic sorting on element composition, the data of five distinct grain size fractions partitioned from four samples of ODP Site 1146 were plotted on Figure 5c. The samples are subparallel to the A-CN line and almost overlapped with the bulk samples from ODP Site 1146, suggesting that the five grain size fractions have similar sediment source as the bulk samples from ODP Site 1146. According to Nesbitt et al. (1996), sorting of materials would result in the compositional changes and thus the plots of distinct grain size fractions would deviate from the predicted weathering trend and form an array oblique but not parallel to the A-CN axis. Obviously, as shown in Figure 5c, the chemical compositions of the sediments at ODP Site 1146 have no evident change resulting from sorting. Thus, neither sea-level changes nor hydraulic sorting had a significant influence on sedimentation at ODP Site 1146 over a million year timescale. Alkali metal elements, such as Na and K, may be mobilized by diagenesis and/or metasomatism.

Fedo et al. (1995) successfully demonstrated a technique for identifying metasomatic effects by use of the A-CN-K diagram. Unmetasomatized sedimentary rocks should plot on the predicted weathering trend that lies parallel to the A-CN axis. If there were considerable addition or substitution of Na and K, then compositions would deviate from the predicted weathering trend and form an array oblique to the A-CN axis (Fedo et al. 1995). The samples analysed from ODP Site 1146 are arrayed subparallel to the A-CN axis. We interpret this to indicate the absence of metasomatism, consistent with their shallow burial (Fig. 5). Sediments from different lithological units generally have different geochemical composition and provenance changes may result in changes in elemental proxies (e.g. Fralick & Kronberg 1997). However, Nd isotopic studies on sediments from ODP Site 1148 indicate that the source of sediment since 23 Ma has been very stable and mainly from South China (Fig. 2) (Clift et al. 2002; Li et al. 2003). The similar mineralogy, grain-size, MAR and geochemistry at ODP Sites 1146 and 1148 are consistent with these two close neighbours having effectively identical provenance since the Early Miocene (Wan et al. 2007). The Pearl River thus appears to be the dominant contributor of detrital material to the northern part of the South China Sea. Concomitant with the onset of the seafloor spreading in the South China Sea (c. 32 Ma; Briais et al. 1993), a large deltaic/estuary system of the palaeo-Pearl River developed on the northern continental margin of the South China Sea (e.g. Guong et al. 1989). In addition, other potentially minor contributions, such as from Taiwan or Luzon and aeolian sources may have supplied some materials to the study area (Shao et al. 2001; Liu et al. 2003; Wan et al. 2007). Grain-size data modelled by an endmember modelling algorithm indicate that a nonnegligible contribution of aeolian dust sourced from arid central Asia to the study site might be expected, especially since about 3 Ma (Wan et al. 2007). The fluvial input also increased rapidly and the proportion of aeolian dust in the total terrigenous content remained constant at that time (Wan et al. 2007). We thus conclude that the Nd isotopes do not require changing sediment sources through that time. We conclude that the selected geochemical proxies did not vary as a result of changes in provenance, but mainly reveal information about chemical weathering conditions and thus the East Asian summer monsoon intensity in South China. As shown in Figure 7, the strikingly similar variation trend of CIA, PIA, Rb/Sr, Ba/Sr, Al/Ti, and K2O/(CaO* þ Na2O) in the terrigenous sediments at ODP Site 1146 shows a general long-term weakening in the degree of chemical weathering

ASIAN SUMMER MONSOON EVOLUTION Fig. 7. Variation of the East Asian summer monsoon intensity since 20 Ma bp, indicated by CIA, PIA, Rb/Sr, Ba/Sr, Al/Ti, K2O/( Na2O þ CaO) at ODP Site 1146. CIA of sediments at ODP Site 1148 (Wei et al. 2006), humidity index from pollen in Sikouzi section (Jiang & Ding 2008), magnetic susceptibility from the Lingtai section on the Loess Plateau (Ding et al. 1999), isotopic composition of black carbon at ODP Site 1148 (Jia et al. 2003), benthic d18O at ODP Site 1148 (Cheng et al. 2004) are also plotted for comparison. See text for explanations. 255

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in South China over the last 20 Ma. Specifically, the chemical weathering intensity, reflected by the combined proxies, is relatively strong before c. 14 Ma, and then slightly weakened between 14 and 8 Ma (Fig. 7). Subsequently, weathering rapidly weakened between about 8.0 and 3.6 Ma and then quickly increased since c. 3.6 Ma. From the point of CIA, PIA, and Rb/Sr ratio, the increase appears not to be permanent and occurs only between c. 3.6 and 2.5 Ma, with a decline resuming since c. 2.5 Ma. In fact, this is not just a local variation limited to the South China Sea because the CIA variations in other three deep drill holes from the Yellow River and Yangtze Delta (Fig. 1), also clearly show a decreasing trend in chemical weathering in the Plio-Pleistocene in eastern China (Yang et al. 2006b). This provides further evidence that the trends are driven by changes in regional environmental conditions.

Evolution of the Asian monsoon system The current climate in South China is mostly controlled by the East Asian monsoon and the precipitation is mainly supplied by the East Asian summer monsoon (Ding 1994). Because continental chemical weathering is largely controlled by moisture and temperature, a wet and warm monsoonal climate will enhance chemical weathering (e.g. Nesbitt & Young 1982; White & Blum 1995). Therefore, variations in chemical weathering intensity in South China revealed by the geochemical records constructed here can essentially be used to reconstruct changes in East Asian summer monsoon intensity (Wei et al. 2004, 2006). Our geochemical data indicate that monsoon intensity gradually decreased after 20 Ma bp and can be roughly subdivided into four stages: a relatively strong monsoon between 20 and 14 Ma, a slight decline during 14 –8 Ma, a rapid weakening during 8.0 –3.6 Ma, and a more variable monsoon since 3.6 Ma bp (Fig. 7). However, considering that decreasing temperatures may result in a decrease in chemical weathering intensity (White & Blum 1995), the global cooling from the Middle Miocene may have at least partly contributed to the observed decrease in chemical weathering intensity in South China. Although it is difficult to clarify the influence of each process on chemical weathering intensity, it is reasonable to assume a dependence of the chemical weathering intensity in South China on the East Asian summer monsoon intensity because of the clear correlation between precipitation and temperature in southern China based on coral record for the last 40 years (Su et al. 2006) and black carbon record since the Early Miocene (Jia et al. 2003; Wei et al. 2006).

Reconstructions of the East Asian summer monsoon since the Early Miocene are very scarce. The magnetic susceptibility in loess-palaeosol sections, a well-known summer monsoon proxy, only spans the base of the Red Clay with an age of about 7–8 Ma (e.g. Ding et al. 1999). As reflected by the loess magnetic susceptibility, the East Asian summer monsoon shows sustained intensification during c. 3.6–2.5 Ma and then become weaker after 2.5 Ma (Ding et al. 1999; An et al. 2001; Sun et al. 2006), which is similar to our geochemical record (Fig. 7). Sedimentation rates at the study site have increased abruptly since about 3 Ma when the summer monsoon became weaker or more variable. In fact, this change is linked to a climate change because in many basins worldwide, increases in sedimentation rates as well in grain sizes of sediments were recorded at c. 2 –4 Ma. Zhang et al. (2001) suggested that disequilibrium states of frequent and abrupt changes in temperature, precipitation and vegetation are effective in causing erosion, but only after 4 –3 Ma, as global climate deteriorated. The significant weakening of the East Asian summer monsoon since the Miocene has also been demonstrated in CIA, Al/Ti, Na/Ti, Ca/Ti, and Al/Na, Al/K, and La/Sm for detrital sediments at ODP Site 1148 in the northern South China Sea (Wei et al. 2006), as well as in the reducing proportion of smectite (a product of chemical weathering) in sediments at the same site (Clift et al. 2002). In addition, evidence for Early to Middle Miocene, especially 17 –15 Ma, strengthening of the East Asian summer monsoon has recently been reported by a much higher humidity index of pollen from the Sikouzi fluviolacustrine sediments on the east side of the Liupan Mountains (Fig. 7) (Jiang & Ding 2008), as well as the lateritic palaeosols in the Nioutoushan basalts, which formed under extraordinary greenhouse climatic conditions with a mean annual temperature of .19 8C and mean annual precipitation of .1650 mm (Zou et al. 2004). The relative decline in sediment flux from East Asia after the maximum during the Early to Middle Miocene (Clift 2006) also supports a reduced summer monsoon intensity since that time. The lowresolution studies based on pollen analysis (Ma et al. 1998; Jiang & Ding 2008) also reveal a general long-term drying trend in the East Asia and thus weakened East Asian summer monsoon intensity since the Miocene. In contrast, such evolution of the East Asian summer monsoon is very different from that seen in the East Asian winter monsoon. Basal dates from the prevailing aeolian Red Clay sediments in the Chinese Loess Plateau indicate initiation of rapid aeolian dust accumulation after about 8 Ma (e.g. Ding et al. 1999; An et al. 2001; Lu et al.

ASIAN SUMMER MONSOON EVOLUTION

2001; Qiang et al. 2001). However, the onset of Asian desertification has been dated to 22 Ma, suggesting a much earlier start to the winter monsoon than previously thought (Guo et al. 2002). Moreover, aeolian sedimentation in the Loess Plateau (Ding et al. 1999; An et al. 2001; Guo et al. 2002) and North Pacific (Rea et al. 1998) increased sharply at about 15 Ma, 8 Ma and 3 Ma (Fig. 8) as well as in West and North China (Fang et al. 1997; Ma et al. 1998; Wang et al. 1999; Sun et al. 2005). These data imply that Asian aridification intensified at 15 Ma, 8 Ma and 3 Ma, possibly accompanied by a strongly intensified East Asian winter monsoon. In the northern South China Sea, clay minerals and grain-size of detrital sediments (Clift et al. 2002; Wan et al. 2007), planktonic foraminifer (P. Wang et al. 2003; Zheng et al. 2004), as well as d13C of black carbon (Jia et al. 2003) strongly indicate that East Asian winter monsoon intensity has generally strengthened since the Early Miocene, especially with three profound shifts at c. 15 Ma, c. 8 Ma and c. 3 Ma (Fig. 8; Wan et al. 2007). In addition to the East Asian summer and winter monsoon, the Asian monsoon system also includes another subsystem, the Indian (or South Asian) summer monsoon. The history of the Indian monsoon is poorly known prior to around 14 Ma, largely because of a lack of suitable core material. Early study of foraminifera and radiolarian assemblages at ODP Site 722 in the Western Arabian Sea/Oman margin revealed that the abundance of G. bulloides and Actinomma sp. (which are associated with the modern summer monsoon) all increased sharply at 8.5 Ma, suggesting that Indian monsoon system was initiated or at least significantly intensified at 8.5 Ma (Kroon et al. 1991; Prell et al. 1992). Studies of the accumulation rates in the Indian foreland and Bengal Fan (Burbank et al. 1993), as well as the Indus Fan (Clift 2006), have suggested reduced erosion following 8.5 Ma. Burbank et al. (1993) suggested that this reduced erosion occurred under the influence of a stronger monsoon. It seems paradoxical because intuitively an increase in rainfall might be expected to increase run-off and erosion, and the Holocene higher rates of erosion in Himalaya are closely linked to periods of stronger summer monsoon (e.g. Goodbred & Kuehl 2000; Clift et al. 2008a). Indeed, Clift (2006) used sedimentation rate data largely from East Asia to argue that the monsoon positively correlates with erosion strength. The highproductivity event identified in the Indian Ocean at 10– 8 Ma was also found in the equatorial Pacific (Gupta et al. 2004) and even throughout the Atlantic (Diester-Haass et al. 2005), implying that the Indian Ocean high-productivity event was not simply the

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result of monsoon-induced upwelling, but may have been caused by strengthened wind regimes resulting from global cooling at that time (Peterson et al. 1992; Gupta et al. 2004). This does not necessarily imply higher rainfall onshore at the same time. We note that a recent study of ODP Site 722 on the Oman margin (Huang et al. 2007) with a higher sampling resolution and more statistically sound estimate compared to Kroon et al. (1991) shows no marked increase in G. bulloides abundance between 10 and 5 Ma. Moreover, the dramatic changes in carbon and hydrogen isotopic ratios of leaf waxes at ODP Site 722 indicate increasing aridity in the continental source regions including Pakistan, Iran, Afghanistan, and the Arabian Peninsula since the Late Miocene (Huang et al. 2007). This supports evidence from the stable isotope ratios of soil carbonates and bivalve shells in Pakistan, India and Nepal indicating a drier climate (Quade et al. 1989, 1995; Dettman et al. 2001; Behrensmeyer et al. 2007). Furthermore, CIA values of sediment recovered from industrial well Indus Marine A1, located offshore the Indus delta (Fig. 1), and which represent chemical weathering in the western Himalayan over the period 17.0 –3.0 Ma (Clift et al. 2008b), also generally decreased since the Middle Miocene, confirming the weakened trend of Indian summer monsoon and concomitant silicate chemical weathering during the Neogene (Fig. 8). Weakening of the Indian summer monsoon and associated drying in the Late Miocene could also explain the reduction in sedimentation rates on the Bengal and Indus Fans at c. 8 Ma and the shift in clay mineralogy observed at 7 Ma (Burbank et al. 1993; Derry & France-Lanord 1996; Clift 2006). Therefore, it appears that the evolution of the East Asian summer monsoon is similar to that of the Indian summer monsoon, but distinct from the East Asian winter monsoon. If that is the case, then the Indian monsoon may also have initiated around the Oligocene–Miocene boundary, as proposed for the East Asian monsoon (Sun & Wang 2005; Clift 2006), although this hypothesis has yet to be tested due to lack of suitable sedimentary records.

Forcing of the Asian monsoons For the past several decades, it is generally accepted that the uplift of the Tibetan Plateau may have played a significant role in strengthening the Asian monsoon and Asian aridification through modulating the atmospheric circulation and its barrier effect to southern-sourced moisture (e.g. Ruddiman & Kutzbach 1989; An et al. 2001; Liu & Yin 2002; Zhang et al. 2007). Although the Tibetan uplift history is far from clear at present (e.g. Molnar

258 S. WAN ET AL. Fig. 8. Marine and terrestrial records since 20 Ma bp showing the evolution of the East Asian summer and winter monsoon, Indian monsoon, global climate, and major uplifts in Tibetan Plateau. The time series are: (a) Rb/Sr (black line) and CIA (grey line) at ODP Site 1146 in the South China Sea; (b) (illite þ chlorite)/smectite (black line) and the coarsest end member EM1 mass accumulation rates (grey line) at ODP Site 1146 (Wan et al. 2007); (c) aeolian MAR from north Pacific ODP Sites 885/886 (Rea et al. 1998) (black line), and mean grain size of quartz at Lingtai Loess section (Sun et al. 2006) (grey line); (d) carbon isotopic ratios of leaf waxes at ODP Site 722 (Huang et al. 2007) (black line) and CIA values at Indus Marine A1 (Clift 2008) (grey line) at the Arabian Sea; (e) Global deep-sea d18O based on data compiled from more than 40 DSDP and ODP sites (Zachos et al. 2001); (f) schematic model showing major uplifts of the Tibetan Plateau (see text references).

ASIAN SUMMER MONSOON EVOLUTION

2005; Harris 2006; Royden et al. 2008), a model of a plateau growing northward progressively through time following the Eocene India-Asia collision seems credible and broadly consistent with the data (Tapponnier et al. 2001). Many dating studies from normal faults (Coleman & Hodges 1995; Blisniuk et al. 2001), north–south oriented dyke swarms in southern Tibet (Williams et al. 2001) and isotope and fossil leaf palaeoaltimetry (Garzione et al. 2000; Rowley et al. 2001; Spicer et al. 2003) suggest that the southern and central Tibetan Plateau has been an established geomorphological feature and possibly remains constant elevation since at least the Middle Miocene (c. 15 Ma) and possibly much older (Rowley & Currie 2006). The evidence for rapid uplift of Tibetan Plateau at about 8 Ma comes from a phase of strong east – west faulting in the Tibetan Plateau (Harrison et al. 1992), from cosmogenic dating of gorge incision in Yunnan starting at 9– 13 Ma (Clark et al. 2005), and from fission-track dating of fault activity of the Liupan Shan thrust (Zheng et al. 2006). The Pliocene uplift of the northern Tibetan Plateau is supported by palaeomagnetic dating of conglomerate and debris-flow deposits at the foot of the Kunlun Mountains at 4.5 –3.5 Ma (e.g. Zheng et al. 2000; Song et al. 2005). Using these constraints we can draw a schematic model (Fig. 8; Wan et al. 2007) as Li et al. (2005) to show the periods of intensified uplift of the Tibetan Plateau in the Late Cenozoic. The stages of strengthening of the Asian winter monsoon correlate well with phases of Tibetan uplift and to Northern Hemisphere Glaciation (Fig. 8; e.g. An et al. 2001; Wan et al. 2007). This study and review shows that the East Asian summer and Indian summer monsoon display a general weakening trend since the Middle Miocene, so that Tibetan uplift may not be the dominant control in forcing the intensity of the Asian summer monsoon. Numerical experiments also suggest that the effect of plateau uplift on the East Asia winter monsoon is more significant than that on the summer monsoon (Liu & Yin 2002), which is consistent with our reconstructions. This finding is very different from the present consensus that Tibetan uplift enhances both the Asian winter and summer monsoons (e.g. An et al. 2001; Zheng et al. 2004). Consequently, we suggest that other mechanisms, for example global cooling, could have had a major effect on the evolution of Asian summer monsoon (Jiang & Ding 2008). From the Early to the Middle Miocene (c. 15 Ma), global ice volume remained low, with a warm phase peaking in the late Middle Miocene climatic optimum (17–15 Ma), and was followed by a gradual cooling and development of a major ice-sheet on

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Antarctica after c. 14 Ma (Flower & Kennett 1994; Zachos et al. 2001). The mean global d18O values then continued to rise gradually through the Late Miocene until the early Pliocene, indicating additional cooling and small-scale ice sheet expansion on West Antarctica (Kennett & Barker 1990) and in the Arctic (Thiede et al. 1998). The late Pliocene is marked by a gradual increase in d18O reflecting the onset of Northern Hemispheric Glaciation (Zachos et al. 2001). In addition, water vapour, Earth’s dominant greenhouse gas, varies in concentration strongly depending on temperature and produces a significant positive feedback in the climate system (e.g. Ruddiman 2002). Therefore, global cooling since the Miocene might be expected to have largely reduced the amount of water vapour held in the atmosphere and thus weakened the Asian summer monsoon, as reflected in the close correlation between the change of Asian summer monsoon intensity and d18O record (Fig. 8).

Conclusions Detailed analysis of major and trace element geochemistry of siliciclastic sediments was carried out on 274 samples from ODP Site 1146 in the northern South China Sea in order to obtain proxy records of the chemical weathering history in South China, which could then be used to assess the evolving strength of the East Asian summer monsoon since 20 Ma bp. From this study, we conclude that geochemical proxies including CIA, PIA, Rb/Sr, Ba/Sr, Al/Ti, and K2O/(CaO* þ Na2O) indicate that the degree of chemical weathering in South China gradually decreased after 20 Ma bp and that this was caused by weakening of the East Asian summer monsoon. The weathering history can be roughly subdivided into four stages: a relatively strong phase between 20 –14 Ma bp, a slight decline at 14 –8 Ma bp, rapid weakening during 8.0 –3.6 Ma bp, and slight recovery with more variability since 3.6 Ma bp. These periods correspond to the periods of the late Middle Miocene climatic optimum, the initiation of Antarctica glaciation, small-scale ice sheet expansion on West Antarctica, and the onset of Northern Hemisphere Glaciation, respectively. We have also shown that the long-term evolution of the East Asian summer monsoon is similar to the Indian summer monsoon but is distinct from the East Asian winter monsoon. While the East Asian summer monsoon has gradually weakened since the Early Miocene the winter monsoon has shown a phased enhancement over the last 20 Ma. The comparison of the long-term evolution of the Asian monsoon and major uplift phases of the Tibetan Plateau strongly suggests that the Asian

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summer and winter monsoon may have different forcing factors. Specifically, the stages in evolution of Asian winter monsoon show a first order correlation to phases of Himalaya-Tibetan Plateau uplift and to Northern Hemisphere Glaciation. However, the correlation between summer monsoon strength and Tibetan uplift is not so good, suggesting that this is influenced by other processes as well. Global cooling since the Miocene may have largely reduced the amount of water vapour held in the atmosphere and thus weakened the Asian summer monsoon. This finding is at odds with the general consensus that Tibetan uplift enhances both the Asian winter and summer monsoons. This research used samples at ODP Site 1146 provided by the Ocean Drilling Program (ODP). ODP is sponsored by the U.S. National Science Foundation (NSF) and participating countries under management of Joint Oceanographic Institutions (JOI), Inc. We thank Dr Sun Youbin, and Kandasamy Selvaraj for kindly providing data of Loess and Taiwan, respectively. We especially thank Editor Ryuji Tada and two anonymous reviewers for thorough and helpful reviews. Thanks also go to Zeng Zhigang, Lu Bo, and Yu Xinke for help during experiments. Funding for this research was supported by the National Natural Science Foundation of China (Grant No. 40706025), Knowledge Innovation Program of CAS (KZCX2-YW-229), National Basic Research Program of China (2007CB411703, 2007CB815903), and China Postdoctoral Science Foundation funded project.

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Colder Subarctic Pacific with larger sea ice caused by closure of the Central American Seaway and its influence on the East Asian monsoon: a climate model study TATSUO MOTOI1* & WING-LE CHAN2 1

Meteorological College, 7-4-81 Asahicho, Kashiwa, Chiba 277-0852, Japan

2

Research Institute for Global Change, Japan Agency for Marine-Earth Science and Technology, 3173-25 Showamachi, Kanazawa-ku, Yokohama, Kanagawa 236-0001, Japan *Corresponding author (e-mail: [email protected]) Abstract: The changes in the sea-ice conditions and sea surface temperatures in the Subarctic Pacific caused by the closure of the Central American Seaway and their influence on the East Asian monsoon are investigated by a series of closed (CE), open (OE) and re-closed (RCE) seaway experiments with a climate model. It is found that a permanent halocline forms in the Subarctic Pacific because of the termination of saline water transport through the seaway in CE and RCE. Efficient cooling by shallow convection in the stratified permanent halocline causes thicker and more extensive sea ice in winter, and leads to colder surface water in summer in the Subarctic Pacific. Colder air, over more extensive sea ice cover in winter and over the colder water in summer, produces higher surface air pressure with anticyclonic wind anomalies in both seasons. Southeasterly and southerly wind anomalies develop around the Japanese archipelago in the East Asian monsoon region and induce warm and humid surface air with increased precipitation over the East Asian continent. These results indicate that the East Asian monsoon is weakened in winter and strengthened in summer as a result of closing the Central American Seaway.

The Asian monsoon influencing the Eurasian Continent and the Indo-Pacific Ocean is one of the most drastic climatic phenomena on the Earth (Clift & Plumb 2008). The evolution of the Asian monsoon and its relationship to tectonic processes are key factors in the natural history of the Earth. The connection between the Asian monsoon and the growth of mountains has been investigated with various types of numerical models, starting from the pioneering work of Manabe & Terpstra (1974). However, the connection between the Asian monsoon and ocean seaways has not been studied, though the Asian monsoon is expected to be strongly influenced by redistribution of sea ice and water temperatures induced by the opening and closing of seaways through changes in ocean-atmosphere interaction. The intensity of the halocline strongly influences sea-ice formation (Carmack 1990), especially in the open deep ocean, because it controls the depth of the convective mixed layer and its heat capacity. As the halocline becomes stronger, the heat capacity of the convective mixed layer becomes smaller because of its shallower depth, which allows sea surface temperatures to reach freezing point and sea ice to form, growing thicker as a result of the usual winter cooling. Tectonic closure

and opening of seaways reorganizes the ocean general circulation and thus influences the development of the halocline, sea surface temperature, sea-ice formation and growth. The Central American Seaway, which was open during most of the Cenozoic and gradually shoaled 16 Ma in the Miocene before final closure at about 3 Ma in the middle Pliocene (Droxler et al. 1998), is thought to have affected the evolution of the Pan-Pacific climate, including the sea surface conditions and the East Asian monsoon. Palaeoclimate proxies suggest that the closure of the seaway caused drastic changes in the Pacific, for example, the onset of stratification with a permanent halocline in the Subarctic Pacific (Sancetta & Silvestri 1986; Haug et al. 1999; Sigman et al. 2004; Swann et al. 2006) accompanied by the establishment of Atlantic-Pacific surface-water salinity contrast (Haug et al. 2001). Proxies also point to larger sea-ice cover in the Bering and Okhotsk Seas with a weakened Kuroshio Current caused by the collapse of North Pacific deep-water formation (Blanc & Duplessy 1982; Sancetta & Silvestri 1986). Several numerical models investigated the effects of the Central American Seaway by using an ocean General Circulation Model (GCM)

From: Clift, P. D., Tada, R. & Zheng, H. (eds) Monsoon Evolution and Tectonics –Climate Linkage in Asia. Geological Society, London, Special Publications, 342, 265–277. DOI: 10.1144/SP342.15 0305-8719/10/$15.00 # The Geological Society of London 2010.

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(Maier-Reimer et al. 1990; Mikolajewicz et al. 1993; Nisancioglu et al. 2003) or a coupled ocean-atmosphere model (Murdock et al. 1997; Mikolajewicz & Crowley 1997; Prange & Schulz 2004; Heydt & Dijkstra 2005; Klocker et al. 2005; Motoi et al. 2005; Schneider & Schmittner 2006; Lunt et al. 2008). Although palaeoclimate studies showed that the opening of the seaway changed the climate in and around the Pacific Ocean drastically, as mentioned above, only Motoi et al. (2005) addressed the changes in the Pacific Ocean. In this present paper, we therefore extend the study by Motoi et al. (2005) to focus on the changes in the Subarctic Pacific sea surface conditions and the East Asian monsoon by using a series of coupled ocean-atmosphere GCM experiments with a closed, open and re-closed Central American Seaway. The results of the experiments are compared to palaeoclimate findings to discuss some of their implications for Asia-Pacific climate evolution. The climate model employed in the present study is of relatively low resolution compared to some recent work, but by using it we can carry out much longer simulations and obtain equilibrium solutions in a short time. The outline of the rest of this paper is as follows. The coupled model and experimental design are described briefly and then the model results for a permanent halocline and ocean conditions within the North Pacific are presented. The influence of the larger sea ice and the colder surface water on the winter and summer monsoons, respectively, is discussed and followed by the palaeoclimatic implications of the present model study and future work. Finally, we summarize our results.

Model and experiments The climate model employed in the present study is a Geophysical Fluid Dynamics Laboratory (GFDL) coupled ocean-atmosphere GCM, as used in Manabe et al. (1991) and Stouffer & Manabe (2003). The resolution of the atmospheric GCM is R15 with nine vertical levels, while that of the ocean GCM is 3.758(lon.)  4.48(lat.) with 12 levels. Because of its low resolution, flux adjustment is necessary for heat and water fluxes to represent realistic climatic sea surface temperature and salinity. Seasonal variation is expressed by the forcing of solar radiation at the top of the atmosphere in the model, estimated from the solar constant and the orbital parameters of the Earth. The model has a simple sea-ice model to calculate sea-ice thickness when the sea surface temperature is below freezing point. The albedo of sea ice depends on surface temperature and ice thickness. The sea ice drifts with the surface ocean currents. The freezing and thawing of sea ice interacts with

the salinity field in the ocean. The oceanic heat flux between ocean and sea ice is expressed reasonably by the conservation of energy. In order to study the influence of the closure of the Central American Seaway on the sea surface conditions in the North Pacific and the East Asian monsoon, three experiments, named Closed Experiment (CE), Open Experiment (OE) and Re-Closed Experiment (RCE), are carried out as shown in Figure 1. The Central American Seaway is closed in CE and RCE and is opened in OE. CE is conducted by using the initial conditions obtained from a GFDL 10 000 year control experiment (Stouffer & Manabe 2003) which were also used in Chan & Motoi (2003, 2005). In OE, four ocean grid boxes are removed from the surface to the 9th level (2559 m) in order to represent the Central American Seaway. The initial conditions of OE are identical to the final state of CE. To close the Central American Seaway again, the removed grid boxes are re-filled in RCE. The final state of the previous OE is used as the initial conditions of RCE. The equilibrium solutions are obtained in each experiment from 5000 year integrations. Figure 2 shows the time series of the North Pacific thermohaline circulation intensity for OE and RCE, defined as the largest value of the meridional overturning stream function. The adjustment time for reorganization of the ocean circulation after opening the Central American Seaway is about 1000 years. The intensity of the North Pacific thermohaline circulation increases from zero to 14 Sv (106 m3s21) during the first 1000 years in OE. It has more or less stabilized after the adjustment and remains at about 14 Sv throughout the last 4000 years, which means that an equilibrium solution has been obtained in OE. In RCE, it decreases from 14 Sv to 6 Sv, with an adjustment time of about 500 years, after which the intensity stabilizes. The data from the last 100 years of each experiment are used in the present study.

CENTRAL AMERICAN SEAWAY

Initial conditions taken from the end of a 10 000year control experiment by Stouffer & Manabe (2003)

Closed Experiment (CE)

5000 a

Open Experiment (OE)

5000 a

Re-Closed Experiment (RCE)

5000 a

TIME

Fig. 1. Diagram illustrating the series of experiments carried out in the present study. The model is integrated for 5000 years in each experiment to obtain equilibrium solutions.

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Fig. 2. Time series of the North Pacific thermohaline circulation intensity for open (OE) and re-closed (RCE) experiments.

Permanent halocline with larger sea ice and colder surface water Figure 3 shows the geographical distribution of current vectors at depths of 170 m for CE, OE and RCE, averaged over the last 100 years of a 5000 year integration. When the Central American Seaway is open, a westward current from the Atlantic to the Pacific Oceans through the seaway develops in the surface layer in OE (Fig. 3b) as a result of a reorganization of the currents. Because the surface water in the Atlantic Ocean is saltier than that in the Pacific, the westward current transports saltier water from the Atlantic to the Pacific in the surface layer. The Atlantic saline water then spreads to the Subarctic North Pacific (40 2 608N) via the North Equatorial Current in the Pacific, the Kuroshio Current, and its extension in the surface layer. The transported salt causes higher sea surface salinity (SSS) in the Subarctic Pacific and its marginal seas in OE. Figure 4 shows the geographical distribution of SSS for CE, OE and RCE. In the Subarctic North Pacific, SSS is lower than 34 psu in CE and RCE (Fig. 4a, c), whereas in OE (Fig. 4b) it is more than 34 psu and about 2 psu higher than those in CE and RCE. In OE the subtropical salinity contrast, which develops in CE and RCE by closure of the Central American Seaway, disappears and the SSSs on both the Pacific and Atlantic sides of the seaway are almost equal at 35 psu. The subsurface salinity distribution in model OE also differs significantly from those in CE and RCE. Figure 5 shows the latitude-depth distribution of zonal-mean salinity in the North Pacific for each experiment. A permanent halocline develops in the Subarctic Pacific and in its marginal seas, as identified in CE and RCE (Fig. 5a, c), while in OE, no such halocline exists and the salinity is vertically uniform with values of about 34.4 psu because of wintertime deep convection (Fig. 5b).

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The salinity distribution is closely associated with the presence and/or absence of deep convection in the Subarctic North Pacific, because water density in high-latitudes is more dependent on salinity. Figure 6 shows the density as a function of salinity and temperature. For high-latitude cold water less than about 10 8C, the contour lines indicate that density depends more strongly on salinity than on temperature. The zonal-mean salinity and temperature at each latitudinal grid-point in the North Pacific with closed, open and re-closed seaway are plotted in this figure. In CE (solid line with black circles) the temperature, salinity and density of the northernmost Subarctic Pacific are 2 8C, 31.45 psu, and 1025.1 kg m23, respectively. When the seaway is open (OE, dashed line with black circles) they are 4.5 8C, 33.7 psu and 1026.7 kg m23. The density becomes larger in OE than in CE because the effect of salinity increase dominates that of the temperature increase. In experiment RCE (dotted line with grey circles), the values are 3 8C, 32.8 psu and 1026.2 kg m23, marking a return to colder, fresher and lighter water. These results indicate that, as a result of higher salinity in OE than in CE and RCE in the Subarctic Pacific, the water becomes heavier, leading to wintertime deep convection and breakdown of the permanent halocline. The deep downwelling of heavier water induces a vigorous thermohaline circulation with a stronger Kuroshio Current, Kuroshio extension and an overall northward component in the interior currents of the shallow layer. Figure 7 shows the stream function of the meridional circulation in the North Pacific for CE, OE and RCE. In OE, as shown in Figure 7b, a vigorous thermohaline circulation reaches water depths of 2000 m with a maximum intensity of 14Sv. In contrast, a permanent halocline forms in the Subarctic Pacific because of the lack of saline water transport through the seaway and the thermohaline circulation is significantly weaker in CE and RCE. Consequently, opening and closure of the Central American Seaway induces significantly different thermohaline circulations in the North Pacific because of differences in the salinity distribution. Figure 8 shows the predicted geographical distribution of sea-ice thickness in February for CE, OE and RCE. Efficient wintertime cooling by shallow convection, caused by stratification of the permanent halocline in CE and RCE, results in thicker and more extensive sea-ice cover in the Okhotsk and Bering Seas, closer to present-day observations. In contrast, in experiment OE, deep convective mixing driven by saline water pumps heat from the deeper layer up to the surface in the Subarctic Pacific, keeping the sea surface temperature there above freezing point throughout the winter, which

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(a)

(b)

(c)

Fig. 3. Geographical distribution of current vectors at depths of 170 m for equilibrium solutions with: (a) closed (CE); (b) open (OE); and (c) re-closed (RCE) seaway. Vector values are averaged over the last 100 years of a 5000 year integration. Arrow scaling is in units of cm s21.

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(a)

(b) (b)

(c)

(c)

Fig. 5. The latitude-depth distribution of zonal-mean salinity in the North Pacific for equilibrium solutions with: (a) closed (CE); (b) open (OE); and (c) re-closed (RCE) seaway. Values are averaged over the last 100 years of a 5000 year integration. Contour interval is 0.2 psu. Fig. 4. Geographical distribution of sea surface salinity for equilibrium solutions with: (a) closed (CE); (b) open (OE); and (c) re-closed (RCE) seaway. Values are averaged over the last 100 years of a 5000 year integration. Contour interval is 1 psu. Regions where salinity is exceeds 34 psu are shaded.

leads to a reduction in sea-ice formation in the Okhotsk and Bering Seas. The seasonal variation in sea-ice thickness averaged over the Okhotsk and Bering Seas for CE (dash), OE (solid) and RCE (dots) are shown in Figure 9. In CE, sea ice exists from November to June and reaches a maximum thickness of 40 cm in March. Sea ice also forms in OE in the Okhotsk and Bering Seas but its maximum thickness is 16 cm, thinner than that in CE. In the case of a re-closed seaway, RCE, a permanent halocline is regenerated in the Subarctic Pacific with shallow wintertime convection and sea ice

thickness of 32 cm, similar to the initially closed seaway case of CE. As shown in Figures 2 and 7, the intensity of the thermohaline circulation in RCE has not diminished completely as in CE because the salinity and density of sea surface water in the Subarctic Pacific halocline are higher in RCE than in CE (see Figs 5 & 6). The findings indicate a bistability in the climate system with/without an active thermohaline circulation in the North Pacific Ocean, associated with a weaker/stronger permanent halocline. This bistability results in sea ice being thinner in RCE than in CE.

Influence of larger sea ice on the winter monsoon Efficient cooling by shallow convection within the permanent halocline induces wintertime formation

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(a)

(b) Fig. 6. Surface density as a function of salinity and temperature. The solid line with black circles, dashed line with black circles and dotted line with grey circles represent zonal-mean salinity and temperature in the North Pacific for equilibrium solutions with a closed (CE), open (OE) and re-closed (RCE) seaway, respectively. Values are averaged over the last 100 years of a 5000 year integration.

of sea ice, which causes significantly cold climate over and around the North Pacific marginal seas driven by its high albedo and barrier effects. In order to clearly detect the influence of the seaway closure, Figure 10 shows the geographical distribution of the differences [closed minus open (CE-OE) and re-closed minus open (RCE-OE)] in surface air temperature during February. The surface air temperatures over the Bering Sea in CE and RCE are 7 and 5 8C lower than those in OE, respectively. The lower surface air temperatures also appear over the Okhotsk Sea and the Subarctic North Pacific. As seen in Figure 7, the thermohaline circulation in CE and RCE is weaker than that in OE, transporting less heat towards the Subarctic North Pacific and its marginal seas. Once there, the transported heat is released to the atmosphere. Thus, the atmosphere over those regions in the CE and RCE experiments receives less heat from the ocean than in OE, resulting in lower surface air temperatures over the Subarctic North Pacific and its marginal seas. The cooler air over the sea ice produces higher surface air pressure with anticyclonic wind anomalies. Figure 11 shows the geographical distribution of the differences [top: closed minus open (CE-OE) and bottom: re-closed minus open (RCE-OE)] in surface air pressure and wind for February. Southeasterly wind anomalies of 1 to 2 ms21 from the North Pacific Ocean to the East Asian Continent, associated with 1 to 5 hPa higher surface air pressure over the thicker and more extensive sea ice in the Okhotsk and Bering Seas,

(c)

Fig. 7. Stream function of the meridional circulation in the North Pacific for equilibrium solutions with (a) closed (CE), (b) open (OE) and (c) re-closed (RCE) seaway. Values are averaged over the last 100 years of a 5000 year integration. Units are in Sverdrups (Sv: 106 m3s21). Contour interval is 2 Sv.

are detected around the Japanese archipelago during the East Asian winter monsoon season. A warmer and more humid climate over the East Asian continent in CE and RCE is induced by the southeasterly wind anomalies. Figure 12 reveals the geographical distribution of the differences [top: closed minus open (CE-OE) and bottom: re-closed minus open (RCE-OE)] in precipitation for February. In CE and RCE, precipitation along the East Asian coast is larger than that in OE by up to 0.2 to 0.4 mm day21. These results indicate

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(a)

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that the East Asian winter monsoon is weakened by closure of the Central American Seaway. The intensity of the atmospheric response to the sea ice depends on its thickness, which is closely related to the strength of the permanent halocline. As seen by comparing the results from experiment CE to those from RCE, when the permanent halocline is stronger, thicker sea ice forms and causes surface air to become cooler. The cooler air in turn causes higher surface air pressure and results in an anticyclonic intensification of the wind fields in the western North Pacific, with a weaker winter monsoon over East Asia.

Influence of colder surface water on the summer monsoon (b)

(c)

Fig. 8. Geographical distribution of sea-ice thickness in February for equilibrium solutions with (a) closed (CE), (b) open (OE) and (c) re-closed (RCE) seaway. Values are averaged over the last 100 years of a 5000 year integration. Contour interval is 1 m.

In summer, sea surface water is colder in experiments CE and RCE compared to OE. Figure 13 shows the geographical distribution of the differences [closed minus open (CE-OE) and re-closed minus open (RCE-OE)], in sea surface temperature in August. The lower sea surface temperatures appear along the North Pacific coast and in the eastern equatorial Pacific region. The sea surface temperatures in the Okhotsk and Bering Seas in CE and RCE are significantly lower by up to 1 –3 8C compared to those in OE. Higher surface air pressure with anticyclonic wind anomalies develops over the North Pacific in experiments CE and RCE, compared to OE. Figure 14 shows the geographical distribution of the differences [top: closed minus open (CE-OE) and bottom: re-closed minus open (RCE-OE)] in surface air pressure and wind in August. Anticyclonic wind anomalies induce offshore Ekman transport anomalies and stronger upwelling in the Subarctic Pacific. The easterly wind anomalies in the equatorial Pacific correspond to La Nin˜a-like lower sea surface temperatures induced by stronger equatorial upwelling. Southerly wind anomalies of about 1 ms21, associated with a surface air pressure ridge over the colder surface water in the Okhotsk Sea, are detected over the Japanese archipelago in the East Asia monsoon region. Warmer and more humid maritime air is transported to the East Asian continent around the Japanese archipelago in CE and RCE by southerly wind anomalies. Figure 15 reveals the geographical distribution of the differences [top: closed minus open (CE-OE) and bottom: re-closed minus open (RCE-OE)] in precipitation in August. In experiments CE and RCE, precipitation over East Asia is larger than that in OE by about 0.5 mm day21. These results indicate that the East Asian summer monsoon is strengthened by closure of the Central American Seaway.

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Palaeoclimatic implications and discussion

Fig. 9. Seasonal variation of sea-ice thickness (cm) averaged over the subarctic Pacific marginal seas. CE: dashed line, OE: solid line, RCE: dotted line.

Fig. 10. Geographical distribution of the differences [top: closed minus open (CE-OE), bottom: re-closed minus open (RCE-OE)] in surface air temperature during February for equilibrium solutions. Values are averaged over the last 100 years of a 5000 year integration. Contour interval is 18C.

The present study examined the results from idealized experiments consisting of a closed, open and re-closed Central American Seaway in a coupled GCM. We focused our attention on the influence of seaway closure and opening on the climate in the Subarctic Pacific and the East Asian monsoon region. By comparing the results of the present model study with those from observed palaeoclimate records, we can discuss some of the implications for the palaeoclimate concerning the closure of the Central American Seaway. One should bear in mind that the present model experiment is limited to modelling the effects of the Central American Seaway and is not designed as a realistic, fully comprehensive palaeoclimate simulation. However, discussions concerning palaeoclimate implications are helpful in understanding the relationship between the tectonically-driven closure of the Central American Seaway and its potential effects on the palaeoclimate evolution of the East Asian monsoon, Subarctic Pacific permanent halocline and sea surface conditions. The present study suggests that surface subtropical saline Atlantic water is transported from the Atlantic Ocean, through the seaway and into the Pacific Ocean and that the sea surface salinity increases in the Subarctic Pacific by about 2 psu. These results are consistent with findings from reconstructed sea surface salinity based on palaeoceanographic proxies from marine sediment cores taken at Ocean Drilling Program (ODP) Site 882 (508220 N, 1678360 E) NW Subarctic Pacific (Swann et al. 2006). They indicate that salinity was 2–4 psu higher in the Subarctic Pacific before closure of the Central American Seaway, around 3 Ma, and that a halocline formed in the North Pacific (Sancetta & Silvestri 1986; Haug et al. 1999; Sigman et al. 2004) with a weakened Kuroshio Current and a collapse of North Pacific deep-water formation (Blanc & Duplessy 1982). A sudden drop in opal accumulation rates has been attributed to the onset of the strong present day-like stratification of the halocline formation in the North Pacific (Sancetta & Silvestri 1986; Haug et al. 1999; Sigman et al. 2004), which might have triggered glaciation over North America (Bartoli et al. 2005; Haug et al. 2005; Swann et al. 2006). The establishment of an Atlantic-Pacific surfacewater salinity contrast (Haug et al. 2001) is also represented in the present model experiments. The model results indicate that thicker and larger sea-ice cover appears in the Bering and Okhotsk Seas following the formation of a permanent halocline caused by the closure of the Central American Seaway. This is consistent with the studies of Blanc & Duplessy (1982) and Sancetta &

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Fig. 11. Geographical distribution of the differences [top: closed minus open (CE-OE), bottom: re-closed minus open (RCE-OE)] in surface air pressure and wind during February for equilibrium solutions. Values are averaged over the last 100 years of a 5000-year integration. Contour interval is 1 hPa. Arrow scaling is in units of m s21.

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(a)

(b)

Fig. 12. Geographical distribution of the differences [top: closed minus open (CE-OE), bottom: re-closed minus open (RCE-OE)] in precipitation during February for equilibrium solutions. Values are averaged over the last 100 years of a 5000 year integration. Contour interval is 0.2 mm day21. Positive values are shaded.

Silvestri (1986). Maslin et al. (1996) noted that a wider continental area covered by glaciers in northeastern Asia might have interacted with the ocean, sea ice and icebergs, which is consistent with the palaeoceanographic finding of an increase in ice-rafted debris at that time (Haug et al. 1999). Features common to both the model results and palaeoclimatic observations encourage us to believe that it will be possible to simulate the past climate more realistically in the future. However, processes in the models, for example, mixed layer schemes, sub-grid scale parameterizations, advection schemes and sea-ice dynamics, need to be improved. Biogeochemical and marine ecosystem processes, which are completely bypassed in this study, are also now in the developing stages,

Fig. 13. Geographical distribution of the differences [top: closed minus open (CE-OE), bottom: re-closed minus open (RCE-OE)] in sea surface temperature during August for equilibrium solutions. Values are averaged over the last 100 years of a 5000 year integration. Contour interval is 0.5 8C.

as mentioned by Schneider & Schmittner (2006). To study the relationship between the ocean seaways and climate evolution in detail by directly comparing simulated model results with observed palaeoclimate records, realistic and high-resolution model experiments with accurate boundary conditions based on sufficient palaeoclimate proxy data are desirable. We also recognize that the seasonality in changes in the East Asian monsoon to the seaway closure is important. In the present study, the monsoon is weakened in winter but is strengthened in summer by the seaway closure. Based on these results, we would recommend that palaeoclimatic proxies for the intensity of the East Asian monsoon be analysed and explained carefully, taking seasonality into consideration.

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Fig. 14. Geographical distribution of the differences [top: closed minus open (CE-OE), bottom: re-closed minus open (RCE-OE)] in surface air pressure and wind during August for equilibrium solutions. Values are averaged over the last 100 years of a 5000 year integration. Contour interval is 1 hPa. Arrow scaling is in units of m s21.

Conclusions Successive experiments consisting of a closed, open and re-closed Central American Seaway show that the opening and closure of the seaway causes a reorganization of the ocean and atmospheric general circulations with significant climate changes in the East Asian monsoon region. When the Central American Seaway is open, a surface westward current, transporting saline water from the Atlantic Ocean to the Pacific Ocean through the seaway, develops from a complete reorganization of the current system. The Atlantic surface

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Fig. 15. Geographical distribution of the differences [top: closed minus open (CE-OE), bottom: re-closed minus open (RCE-OE)] in precipitation during August for equilibrium solutions. Values are averaged over the last 100 years of a 5000 year integration. Contour interval is 0.5 mm day21. Positive values are shaded.

saline water is carried by the anticyclonic gyre circulation of the North Equatorial Current, the Kuroshio Current, and its extension, spreading into the Subarctic Pacific Ocean. This causes the sea surface salinity in the Subarctic Pacific and its marginal seas to become about 2 psu higher in experiment OE compared to in CE and RCE, and the salinity contrast between the subtropical Pacific and Atlantic Oceans to diminish. A permanent halocline forms in the Subarctic Pacific in the CE model and reappears in RCE with thicker and more extensive sea-ice cover and a weaker thermohaline circulation. However, in the OE model, the halocline is destabilized by deep convection as a result of saltier and denser surface water. Sea-ice cover retreats and becomes

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thinner, while the thermohaline circulation intensifies with a stronger Kuroshio Current and Kuroshio extension into the North Pacific surface layer. Efficient wintertime cooling by shallow convection, driven by stratification of the permanent halocline in CE and RCE models, enhances sea-ice formation which causes significantly colder air over and around the Okhotsk and Bering Seas. In contrast, in the OE model, the wintertime sea surface temperature in the Subarctic Pacific is warmer because deep convection, induced by saline water, pumps heat from the deeper layer up to the surface, leading to reduced sea-ice formation, causing warm and humid climate over the Okhotsk and Bering Seas. In the CE and RCE models, colder air over the thicker and more extensive sea ice caused by the closure of the seaway produces higher surface air pressure with southeasterly wind anomalies in the East Asian winter monsoon region. The southeasterly wind anomalies lead to warm and humid air over the East Asian continent, indicating that the East Asian winter monsoon is weakened by closure of the seaway. The intensity of the atmospheric response depends on sea-ice thickness, which is controlled by the strength of the permanent halocline and thermohaline circulation. In the case of a stronger permanent halocline with weaker thermohaline circulation, thicker sea ice forms and induces colder surface air with higher surface air pressure and anticyclonic intensification of wind anomalies over the Subarctic Pacific, Okhotsk and Bering Seas, resulting in weaker winter monsoon in the East Asian region. In summer, sea surface water along the North Pacific coast and in the eastern equatorial Pacific region is colder in models CE and RCE than in OE. The higher surface air pressure with anticyclonic wind anomalies, which develop over the North Pacific in models CE and RCE, induces stronger upwelling along the North Pacific coast and in the equatorial Pacific. Much lower sea surface temperatures in models CE and RCE are produced in the Bering and Okhotsk Seas because of a permanent halocline in the colder North Pacific Ocean. Southerly wind anomalies develop in the surface air pressure ridge over the significantly colder surface water in the Okhotsk Sea. These wind anomalies carry warmer and more humid maritime air to the East Asian continent around the Japanese archipelago, resulting in a stronger summer monsoon over East Asia. The authors wish to thank S. Manabe and R.J. Stouffer for helping to set up the GFDL model and for their invaluable advice and suggestions. The manuscript also benefited greatly from comments by the editor and reviewers. This study was partially supported by the Grant-in-Aid

for Scientific Research (C20540433) from the Ministry of Education, Culture, Sports, Science and Technology, Japan.

References Bartoli, G., Sarnthein, M., Weinelt, M., Erlenkeuser, H., Garbe-Schonberg, D. & Lea, W. 2005. Final closure of Panama and the onset of Northern Hemisphere glaciation. Earth and Planetary Science Letters, 237, 33– 44. Blanc, P. L. & Duplessy, J. C. 1982. The deepwater circulation during the Neogene and the impact of the Messinian salinity crisis. Deep-Sea Research, 29, 1391– 1414. Carmack, E. C. 1990. Large-scale physical oceanography of polar oceans. In: Smith, W. O., Jr. (ed.) Polar Oceanography. Part A: Physical Science. Academic Press, San Diego, 171– 222. Chan, W.-L. & Motoi, T. 2003. Effects of stopping the Mediterranean outflow on the southern polar region. Polar Meteorology and Glaciology, 17, 25–35. Chan, W.-L. & Motoi, T. 2005. Response of thermohaline circulation and thermal structure to removal of ice sheets and high atmospheric CO2 concentration. Geophysical Research Letters, 32, doi: 10.1029/2004GL021951. Clift, P. D. & Plumb, R. A. 2008. The Asian Monsoon: Causes, History and Effects. Cambridge University Press, Cambridge. Droxler, A. W., Burke, K. C. et al. 1998. Caribbean constraints on circulation between Altantic and Pacific oceans over the past 40 million years. In: Crowley, T. J. & Burke, K. C. (eds) Tectonic Boundary Conditions for Climate Reconstructions. Oxford University Press, New York, 169– 191. Haug, G. H., Sigman, D. M., Tiedemann, R., Pedersen, T. F. & Sarnthein, M. 1999. Onset of permanent stratification in the subarctic Pacific Ocean. Nature, 401, 779 –782. Haug, G. H., Tiedemann, R., Zahn, R. & Ravelo, A. C. 2001. Role of Panama uplift on oceanic freshwater balance. Geology, 29, 207 –210. Haug, G. H., Ganopolski, A. et al. 2005. North Pacific seasonality and the glaciation of North America 2.7 million years ago. Nature, 433, 821– 825. Heydt, A. von der & Dijkstra, H. A. 2005. Flow reorganizations in the Panama Seaway: a cause for the demise of Miocene corals? Geophysical Research Letters, 32, doi: 10.1029/2004GL020990. Klocker, A., Prange, M. & Schulz, M. 2005. Testing the influence of the Central American Seaway on orbitally forced Northern Hemisphere glaciation. Geophysical Research Letters, 32, doi: 10.1029/ 2004GL021564. Lunt, D. J., Valdes, P. J., Haywood, A. & Rutt, I. 2008. Closure of the Panama Seaway during the Pliocene: implications for climate and Northern Hemisphere glaciation. Climate Dynamics, 30, 1– 18. Maier-Reimer, E., Mikolajewicz, U. T. & Crowley, J. 1990. Ocean general circulation model sensitivity experiment with an open central American isthmus. Paleoceanography, 5, 349–366.

COOLING THE NORTH PACIFIC BY GATEWAY CLOSURE Manabe, S. & Terpstra, T. B. 1974. The effects of mountains on the general circulation of the atmosphere as identified by numerical experiments. Journal of the Atmospheric Sciences, 31, 3– 42. Manabe, S., Stouffer, R. J., Spelman, M. J. & Bryan, K. 1991. Transient response of a coupled oceanatmosphere model to a gradual change of atmospheric CO2. Part I: annual response. Journal of Climate, 4, 785–818. Maslin, M. A., Haug, G. H., Sarnthein, M. & Tiedemann, R. 1996. The progressive intensification of northern hemisphere glaciation as seen from the North Pacific. Geologische Rundschau, 85, 452–465. Mikolajewicz, U. & Crowley, T. J. 1997. Response of a coupled ocean/energy balance model to restricted flow through the central American isthmus. Paleoceanography, 12, 429– 441. Mikolajewicz, U., Maier-Reimer, E., Crowley, T. J. & Kimm, K. Y. 1993. Effect of Drake and Panamanian gateways on the circulation of an ocean model. Paleoceanography, 8, 409 –426. Motoi, T., Chan, W.-L., Minobe, S. & Sumata, H. 2005. North Pacific halocline and cold climate induced by Panamanian Gateway closure in a coupled oceanatmosphere GCM. Geophysical Research Letters, 32, doi: 10.1029/2005GL022844. Murdock, T. Q., Weaver, A. J. & Fanning, A. F. 1997. Paleoclimatic response of the closing of the Isthmus of Panama in a coupled ocean-atmosphere model. Geophysical Research Letters, 24, 253– 256.

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The impacts of Tibetan uplift on palaeoclimate proxies DANIEL J. LUNT1,2*, RACHEL FLECKER1 & PETER D. CLIFT3 1

BRIDGE, School of Geographical Sciences, University of Bristol, University Road, Bristol BS8 1SS, UK

2

British Antarctic Survey, Geological Sciences Division, High Cross, Madingley Road, Cambridge CB3 0ET, UK 3

School of Geosciences, Meston Building, King’s College, University of Aberdeen, Aberdeen AB24 3UE, UK *Corresponding author (e-mail: [email protected]) Abstract: Several palaeoclimate proxy records have been interpreted as representing the direct effects of Tibetan uplift on climate, and particularly the intensity of the Asian summer monsoon. However, there are other possible causes for the transitions and changes which have been observed, such as varying greenhouse gas concentrations, nodes or extremes in orbital forcing, and changing continental configurations. In this study we model the direct effects of Tibetan uplift on sea surface temperatures (SSTs), vegetation, and river discharge. We investigate whether these climatic effects of topographic uplift are likely to be detectable in proxy records, and also whether the proxies could be used to distinguish between different paradigms for the history of plateau uplift. We find that the SSTs in the western Pacific, South China Sea and Indian Ocean are generally insensitive to Tibetan uplift; however, vegetation in the region of the plateau itself, and river discharge from the Yangtze, Pearl, and in particular the Ganges/ Brahmaputra, could provide a good test of our understanding of Tibetan uplift history.

A key contributor to the characteristics of the East Asian Monsoon (EAM) is the seasonal temperature range of the Tibetan Plateau (e.g. Kutzbach et al. 1989). As a result, the long-term evolving intensity of the EAM is thought to be intrinsically linked to the nature and timing of the tectonic uplift of Tibet (An et al. 2001; Kitoh 2004). However, the wide range of proxy records used to monitor the EAM (e.g. loess, marine, estuarine and lake sediments, ice cores, speleothems and tree rings) are often contradictory, so that estimates of monsoon timing and intensity (both Quaternary and preQuaternary) vary considerably (e.g. Raymo & Ruddiman 1992; An et al. 2001; Thamban et al. 2002; Chen et al. 2003; Clift et al. 2004; Liu et al. 2004). An alternative approach initially developed by Hahn and Manabe (1975) is to use well designed experiments with numerical models to evaluate the direct impact of Tibetan uplift on the Asian Monsoon system (e.g. Prell & Kutzbach 1992). In this paper we document the results of simulating Tibetan uplift, and examine two different uplift histories for Tibet. Specifically, focusing on SST, vegetation, and run-off, we consider whether locations previously used to generate proxy records for evaluating past monsoonal variability are those most sensitive to changes induced by Tibetan uplift.

In general, previous modelling studies have used very similar methodology: typically, a series of simulations have been carried out with linearly increasing topographic height, either in the region of the Tibetan Plateau alone (e.g. Liu & Yin 2002) or globally (e.g. Kitoh 2004). This experimental design is useful for understanding the mechanisms by which orography influences monsoon systems, but is at odds with more recent data that support earlier models that the Tibetan Plateau likely uplifted with orogeny occurring first in southern Tibet (England & Searle 1986), and spreading northwards and eastwards with time (Clark et al. 2005; Rowley & Garzione 2007). As such, some of the conclusions that such modelling studies reached, for example, that the onset of the monsoon occurred when the Plateau reached 50% of its current height (Kitoh 2004), are possibly misleading because this particular configuration may never actually have existed. The study of Ramstein et al. (1997) simulated more realistic palaeogeographies, but used prescribed rather than computed SSTs, and, as a result, missed the important amplification of the tectonic-driven climate response (Kitoh 2004), particularly via western Pacific SST changes. Indeed, most previous studies have been carried out with prescribed SSTs (e.g. An et al. 2001; Liu & Yin 2002). Those which have utilized

From: Clift, P. D., Tada, R. & Zheng, H. (eds) Monsoon Evolution and Tectonics –Climate Linkage in Asia. Geological Society, London, Special Publications, 342, 279–291. DOI: 10.1144/SP342.16 0305-8719/10/$15.00 # The Geological Society of London 2010.

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fully-coupled atmosphere-ocean models with computed SSTs (e.g. Abe et al. 2005) have carried out idealized simulations, lowering orography globally. Many of the numerical simulations have been integrated for a relatively short time, typically 10 years (e.g. Zhongshi et al. 2007). While this may allow equilibrium to be reached in an atmosphere-only simulation, the chaotic nature of the climate system means that some of the published studies may be influenced by interannual variability, and consequently some of the conclusions reached may have been erroneous. The fully-coupled atmosphere-ocean study of Kitoh (2004) was integrated for 50 years, but this is still some way off an equilibrium state in the intermediate ocean, let alone the deep ocean. In summary, although there have been interesting modelling studies carried out to date, some of the problems inherent in previous work can now be addressed. Namely: fully-coupled atmosphereocean General Circulation Model (GCM) simulations representing more recent ideas concerning topographic history of the Tibetan Plateau, carried out over a meaningful length of time to allow equilibrium to be reached. We describe the models used in this study and the experimental design. We then present the results and their implications for reconstructing Tibetan uplift from proxies. Finally, we compare our results with those of previous studies, and discuss the limitations of our models and approach and present our conclusions.

Model descriptions We make use of the UK Met Office GCM, HadCM3L (Cox et al. 2001). The horizontal resolution of the atmospheric model is 2.58 in latitude by 3.758 in longitude, with 19 vertical layers. The atmospheric model has a time step of 30 minutes and includes many parameterizations, representing subgridscale effects, such as convection (Gregory & Rowntree 1990), boundary layer mixing (Smith 1993), and radiation (Edwards & Slingo 1996). The land-surface scheme (MOSES 2.1) contains a representation of evaporation, which includes the dependence of stomatal resistance on temperature, vapour pressure and CO2 concentration (Cox et al. 2001). The spatial resolution in the HadCM3L ocean is 2.58 in latitude by 3.758 in longitude, with 20 vertical layers. The model uses the Gent and McWilliams (1990) mixing scheme, and there is no explicit horizontal tracer diffusion. The sea-ice model (Hibler 1979) uses a simple thermodynamic scheme and contains parameterizations of ice drift and leads (Cattle & Crossley 1995). Compared to the original Cox et al. (2001) study,

our model has no need for an oceanic flux correction to maintain an ‘on’ state of the thermohaline circulation in the modern control simulation, because of the replacement of Iceland in the land-sea mask with ocean bathymetry typical of the North Atlantic. This technique is also used in the FAMOUS GCM (Smith et al. 2008), which is identical to HadCM3L except that it uses a lower resolution atmosphere. HadCM3, which is identical to HadCM3L except that it uses a higher resolution ocean and has an older version of the land-surface model (MOSES-1), is described in some detail, along with an assessment of the climatologies it produces under modern boundary conditions, in Gordon et al. (2000). Both future and palaeoclimate predictions from HadCM3 have been reported extensively in the most recent report of the Intergovernmental Panel on Climate Change (IPCC, Solomon et al. 2007), and in terms of its simulation of key aspects of current climate, such as surface air temperature, it performs well compared to other GCMs of similar complexity (Covey et al. 2003). HadCM3L has been successfully used in previous studies of pre-Quaternary climates (e.g. Lunt et al. 2007). In a study such as this, it is essential to verify that the tools are appropriate for the task to be undertaken. In this case, it is important that HadCM3L has a satisfactory representation of the current state of East Asian climate, and in particular the Asian monsoon system; otherwise, its predictions for palaeomonsoons would be somewhat dubious. Jiang et al. (2005) looked explicitly at the simulation of the East Asian climatology in seven fully coupled GCMs, including HadCM3. These authors found that the models could successfully reproduce the annual and seasonal surface air temperature and precipitation climatology in East Asia. They found that although no single model performed best according to every metric considered, HadCM3, along with ECHAM4 (Roeckner et al. 1996), performed much better than the other GCMs, and better than the model ensemble mean. Of particular interest for this study is the summer precipitation, for which HadCM3 performed best of all the GCMs considered. Turner et al. (2005) examined boreal summer tropical SSTs, precipitation, and 850 mbar winds from HadCM3 and also found that the mean state of the monsoon was well simulated. However, they found that the teleconnection between ENSO and the monsoon was rather poorly represented. Inness and Slingo (2003) made a similar comparison for boreal winter, autumn and spring, and reached similar conclusions. In summary, we consider that HadCM3L is an appropriate tool for use in this study, in particular given its performance relative to other models of similar complexity, especially with respect to precipitation,

TIBETAN UPLIFT AND PALAEOCLIMATE PROXIES

although with proviso that care must be taken if it is used for interpreting ENSO-related signals. We also make use of the BIOME4 vegetation model (Kaplan 2001). BIOME4 employs a detailed vegetation classification scheme by discriminating between 28 potential natural vegetation types. BIOME4 is used to translate the multivariate climate data from our GCM simulations, specifically the seasonal cycle of temperature, precipitation and solar radiation, into biome distributions. BIOME4 is a mechanistically based model that was developed from physiological considerations that place constraints on the growth and regeneration of different plant functional types (e.g. cool conifer forest, tropical grassland). These constraints are calculated through the use of limiting factors for plant growth, which include the mean temperature of the coldest and warmest months, the number of growing degree days (GDD) above 0 and 5 8C (i.e. the total number of days during which the threshold temperature is passed, multiplied by the average amount in degrees by which the threshold is passed on those days), and the calculation of a coefficient (Priestley-Taylor coefficient) for the extent to which soil moisture supply satisfies atmospheric moisture demand. GDDs are calculated by linear interpolation between mid-months, and by a one-layer soil moisture balance model independent of the HadCM3L GCM hydrology (Prentice et al. 1992, 1993). We run BIOME4 on a 18  18 grid. Monthly mean surface temperature, precipitation and cloud cover were compiled from our uplift simulations to provide the climatic information necessary for the BIOME 4 model. When forcing biomes, we used a standard anomaly method, as employed in several previous studies (e.g. Haxeltine & Prentice 1996; Texier et al. 1997), whereby a correction factor derived from the GCM systematic error relative to present-day observations is applied to the GCM uplift climatologies. BIOME4 has been used to simulate vegetation patterns and vegetation-climate feedbacks for the Quaternary (e.g. Claussen et al. 1998; Prentice et al. 1998). However, Haywood et al. (2002), Francis et al. (2007) and Salzmann et al. (2008) have demonstrated that BIOME4 can also be successfully applied to global model simulations of pre-Quaternary time intervals that include modified orography.

Experimental design In order to assess the impact of Tibetan uplift on palaeoclimate proxies, and the importance of uplift history, we have carried out four simulations with HadCM3L, which represent idealized (a) preuplift, (d) post-uplift (i.e. present-day), and (b,c)

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syn-uplift configurations. The pre-uplift configuration (Eflat) represents a world in which all orography is the same as that of modern (Emod) except in the region of Central Asia, in which the orography is set flat to 1000 m (similar to the surrounding regions). The two syn-uplift configurations represent (b) the more traditional syn-uplift configuration (Efrac), in which all orography in the Tibetan region is uplifted by a constant fraction (45%) from the pre-uplift case towards modern, and (c) a syn-uplift configuration (Esouth) following Rowley and Garzione (2007), in which the Tibetan Plateau grows from the south, being equal to Emod south of 358N and equal to Eflat north of 358N. The fractional uplift in (b) is chosen so that the mean orography in the Tibetan region is the same as that in (c). The orography in Central Asia in the model in these four cases is shown in Figure 1. The orography outside SE Asia, and all other boundary conditions in the GCM (e.g. land-sea mask, atmospheric CO2, vegetation, soils, ice sheets) are representative of modern pre-industrial (c. 1850) conditions. The syn-uplift configurations are not strongly constrained by geological evidence, but are instead intended to represent two end-members of possible uplift histories, as a basis for this study. The four GCM simulations were each run for a total of 125 years, and the results discussed in this paper are means of the last 35 years. Although this is shorter than typical timescales of ocean circulation, the four simulations have all reached a quasi-equilibrium state, at least in terms of the surface variables, at the end of the spin-up period, the trend in 2 m air temperature over the last 100 years of each simulation is less than 0.1 8C per century. It is possible that, given longer, the ocean could respond in a non-linear fashion and lead to either increased or reduced sensitivity to the orographic forcing. However, our approach has also been used by Kitoh (2004), who integrated for a lower total of 50 years.

Results In this section we present and discuss the results of the four simulations. We focus on annual mean SST, vegetation, and run-off changes, as these are reconstructed by palaeoclimate proxies of the monsoon (e.g. through alkenones, oxygen isotopes, pollen, C/N ratios, sea surface salinity (SSS) etc.). They are also relatively well simulated by the model for modern conditions. The model control simulation (pre-industrial) has annual mean SSTs in very good agreement with the HadISST observational dataset (Rayner et al. 2003) in the Indian Ocean, but there is a cold bias in the north-western

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(b)

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during−uplift (fractional )

pre−uplift

(d)

during−uplift (from south)

Fig. 1. East Asian orographic boundary condition [m] applied in the 4 simulations in this study (a) Eflat, (b) Efrac, (c) Esouth, and (d) Emod. Black arrows show the direction of geological time. The mean orography over the East Asian region in configurations (b) Efrac and (c) Esouth is the same.

D. J. LUNT ET AL.

(c)

post−uplift

TIBETAN UPLIFT AND PALAEOCLIMATE PROXIES

Pacific of around 4 8C. The simulation of modern natural vegetation by the BIOME model is in good agreement with data in East Asia (Haxeltine & Prentice 1996), although the extent of the Gobi and Taklimakan deserts is too great. Comparison of model results with run-off data is problematic as current observations are likely highly influenced by the presence of dams and other anthropogenic effects. However, the control model precipitation, which is the main driver of the run-off signal, is relatively good in East Asia compared to other models of similar complexity (Jiang et al. 2005; Solomon et al. 2007). We also discuss the mechanisms behind the observed changes. Our aim is to investigate the effect of Tibetan uplift on measurable palaeoclimate proxies, and to explore the importance of uplift history for interpreting these changes. Our basic assumption is that if there is an observable difference between the Eflat and Emod cases then the proxy is a good indicator for climate change caused by uplift, and if there is an observable difference between the Esouth and Efrac cases then the proxy is a good indicator for distinguishing between climate change caused by contrasting uplift histories.

Sea surface temperature The changes in SST in the uplift simulations (Efrac, Esouth, Emod), relative to the flat case (Eflat), are shown in Figure 2. The primary response in all three uplift simulations is a cooling in the north-western Pacific, of the order of 1.5 8C in the Emod case (Fig. 2c). The temperature signal is strongest in the boreal winter (December/January/ February–DJF). This SST cooling is related to the temperature response to the uplift on land (Fig. 3a). The main continental change is unsurprisingly a lapse-rate effect over the uplifted region itself. This is accompanied by a tripole pattern of year-round warming to the north of the plateau, and cooling of the Pacific Ocean. The warming to the north of the uplifted region is also seen in the simulations of Abe et al. (2005), which is attributable to low precipitation rates and low soil moisture, presumably inhibiting latent cooling. In our runs, this drying north of the plateau is also observed (Fig. 3b), although the drying is greater during the boreal summer (June/July/August– JJA) than in DJF. The larger temperature response in DJF can be attributed in our simulations to the fact that there is a marked increase in albedo due to enhanced snow cover following uplift, leading to positive feedback. The temperature changes are related to the changes in atmospheric circulation. The uplift of Tibet leads to an intensification of the main pressure systems associated with the East Asian monsoon that is, the western Pacific

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subtropical high gets higher with uplift, and the monsoon low gets lower. Figure 2a, b show the SST change for the two uplift histories, Efrac– Eflat and Esouth–Eflat. The main temperature changes are remarkably similar for the two alternative syn-uplift cases. This indicates that the uplift history is not particularly important for understanding the large-scale SST effects of Tibetan uplift. However, in terms of detecting a signal due to uplift in SST proxies (Fig. 2c), there are local regions where there are distinct differences in SST (e.g. in the Sea of Okhotsk, where there is also an increase in winter sea-ice associated with the cooler temperatures following uplift). However, at these local scales the GCM results are not robust and are likely to be model-dependent (see Discussion). The exact history of uplift also has little influence on the pressure patterns (e.g. 850 mbar wind field), either annually or seasonally. Therefore, it appears that SST is not a good indicator for climate change caused by Tibetan uplift or the nature of that uplift history. There are isolated regions which would show a signal, but these would be hard to interpret because of their relatively small magnitude and limited spatial extent, making any observed changes very hard to attribute to uplift directly. In addition, no change in SSTs is predicted as a result of Tibetan uplift and associated monsoonal variability in localities such as the South China Sea, a classic region for monsoon studies and where Holocene changes to monsoon intensity have been documented (e.g. Huang et al. 1997).

Vegetation Figure 4 shows the vegetation distributions for the four simulations, as predicted by the BIOME4 model. Here, the most marked changes in vegetation are over the uplifted region itself. In the Eflat case (Fig. 4a), the plateau is essentially bare soil in the west, and shrubland in the east. After uplift (Fig. 4b), the southern plateau includes various tundra biomes, as well as scattered forests. Locally, the uplift history can be nicely diagnosed from the various vegetation types. For example, in the Esouth case (Fig. 4d), the vegetation in the uplifted region in the south is almost identical to the fully uplifted Emod case. The strong dependence of local vegetation on the uplift is not surprising given that locally the temperature changes are large, mainly due to the lapse-rate effect (Fig. 3a). However, outside the plateau region, the vegetation distribution is relatively unchanged in the different simulations. There are some meridional shifts in biome boundaries, such as between warm mixed and temperate deciduous forests in eastern China, but these are likely to be model-dependent, and exist only over a relatively narrow band of latitudes.

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(a)

(b)

(c)

Fig. 2. (a,b,c) Annual mean SST change [8C] (a) Efrac– Eflat, (b) Esouth–Eflat, and (c) Emod–Eflat. Solid contours represent positive values; dotted contours represent negative values.

The vegetation distribution is controlled mainly by temperature and precipitation, because the far-field temperature and precipitation changes are of relatively small magnitude (Fig. 3a, b), it is unsurprising that they would not be detected by a vegetationbased palaeoclimatic proxy. The temperature and precipitation changes are broadly consistent with those by previous workers who have investigated the impacts of uplift on East Asian climate, but the precipitation change may be model dependent (see Discussion). The largest precipitation signal (a)

we see is a wettening in the continental region of uplift in JJA, and a dipole in the tropical western Pacific of wettening to the north and drying to the south. In addition, there is a wettening in eastern China in DJF, and a drying in western India in JJA (which is visible in the annual mean anomaly, see Fig. 3b). The increased precipitation in JJA over the plateau is most likely orographically induced; as the moist air comes off the Indian Ocean in JJA it rises when it meets the plateau, and saturates as it cools. Over continental SE Asia there is little (b)

Fig. 3. (a) Annual mean surface air temperature change [8C], Emod– Eflat. (b) Annual mean precipitation change [mm day21], Emod– Eflat. Solid contours represent positive values; dotted contours represent negative values.

TIBETAN UPLIFT AND PALAEOCLIMATE PROXIES

Fig. 4. Vegetarian biome types predicted by the BIOME4 model for the four simulations: (a) Eflat; (b) Efrac; (c) Esouth; and (d) Emod.

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difference between the two histories in DJF. In contrast, over the western tropical Pacific there is a large difference; namely, that the dipole in precipitation is only present in the Esouth case. In JJA, the opposite is the case, because the precipitation changes over land are more intense in the Efrac case than the Esouth case, in particular the drying in western India, and the wettening in eastern Tibet. We conclude that vegetation changes in the region of the Tibetan Plateau should be a very good indicator of uplift and the uplift history. Outside of the plateau itself, the vegetation signal associated with uplift is very weak. There are some latitudinal shifts in vegetation type, but these are relatively small and would be almost impossible to ascribe to uplift as opposed to other possible forcing factors.

Run-off From the model results it is possible to estimate the change in run-off from various river basins due to the uplift of Tibet. The model calculates run-off based on a ‘leaky bucket’ parameterization, in which any falling continental precipitation is partitioned between immediate ‘fast’ run-off, and a component which is combined with evaporation to determine the soil moisture. If the soil moisture saturates, then additional precipitation leads to ‘slow’ run-off. All run-off is transferred immediately to the ocean, via an assigned river mouth. Unfortunately, the default river basins in the model do not actually correspond very well with reality in SE Asia. Because of this, we have improved the definitions for the purposes of this study. We assume that the river basins have not changed dramatically as a direct effect of uplift (see next section for further discussion), despite suggestions from some authors that drainage has experienced major re-organization (Clark et al. 2004; Brookfield 1998). Our improved river routing, the water flux for each river, and the total outflow at the major river mouths of East Asia is shown in Figure 5. Here, the uplift of Tibet does have a significant impact on the outflow from the Ganges/Brahmaputra, the Yangtze, and the Pearl rivers (Fig. 5e, comparison of Emod with Eflat for each river system). However, the other major rivers, the Amur, Yellow, Mekong, Salween, and Irrawaddy Rivers, do not show significant changes in outflow. The uplift history (Fig. 5(e), comparison of Efrac with Esouth for each river system) does have an important role in determining the run-off into the Ganges/Brahmaputra (although not so much for the Yangtze or Pearl). Specifically, the fractional uplift increases Ganges outflow relative to the flat case by about a factor of 3, the uplift to the south by a factor of 5, and total uplift by a factor of 7.

These large magnitude changes in discharge should be identifiable in a variety of different proxy records through their impact on the freshwater flux to estuaries and marginal marine settings. The proxies used to monitor fluctuations in the freshwater flux include carbon isotopes, C/N ratios, Branched and Isoprenoid Tetraether (BIT) index (Hopmans et al. 2004), and sea surface salinity (SSS) as reconstructed by paired measurements of d18O and alkenones. Studies reconstructing the freshwater flux of rivers in SE Asia by focusing on the estuarine settings are limited, but a successful project to document the monsoonal impact on the Pearl River during the Holocene has been carried out (Zong et al. 2006) and this approach could be applied to older sediments providing the catchment of the river has remained relatively constant. The more established method of monitoring variations in the freshwater flux through SSS reconstructions has been successfully applied to the South China Sea (Wang et al. 1999). The South China Sea is not the ideal place to look for records of monsoonal variability induced by changes in Tibetan altitude because, considering the major rivers feeding this semienclosed basin, whereas the Pearl’s discharge varies markedly, the Mekong has a much more constant discharge (Fig. 4e). As a result the sort of changes in surface salinity (SSS) documented by Wang et al. (1999) for the Late Holocene are unlikely to produce useful records for testing the different Tibetan uplift histories. Therefore, run-off can be used a good indicator of both uplift and uplift history, provided that the correct river basin is examined. Specifically, our results indicate that discharge from the Ganges/ Brahmaputra would be the best indicator of uplift history, and that the Pearl and Yangtze would also show significant uplift signals.

Discussion Climate models are valuable tools for understanding past climate change, and interpreting proxy records. However, models are by their very nature simplified representations of the real world, and results may be dependent on the model, or version of the model, used. As such, modelling studies should always be compared with similar experiments carried out with other models, and care must be taken when interpreting their results. Here we compare with previous works, discuss the caveats and limitations of our simulations, and highlight some directions for future work. Comparison with previous work is hampered by the fact that no one has carried out identical simulations to the ones we present here with a model of

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Fig. 5. (a– d) Major river run-off in the four simulations. The straight lines join the model continental grid cells with their corresponding river mouth. The thickness of the straight lines indicates the amount of run-off from that grid cell. The area of the filled circles at each river mouth represents the total run-off which enters the ocean at that point. Only the major rivers are shown. (a) Eflat, (b) Efrac, (c) Esouth and (d) Emod. (e) shows the total outflow for the major rivers, normalized so that each river is given a value of 1 for the Eflat case -for each river from left to right the bars represent simulations Eflat, Efrac, Esouth and Emod.

similar complexity. However, similar experiments with various models can be compared. Kitoh (2004) and Abe et al. (2003) analyse a set of experiments in which global topography is uplifted

incrementally from a flat Earth towards modern, in the MRI coupled atmosphere-ocean model. They present the change in summer SST between their modern and flat Earth simulations [Abe et al.

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(2003), their fig. 14f, and Kitoh (2004), their fig. 13]. Our SST results for the Emod– Eflat case are remarkably similar to those of Abe et al. (2003) in the Western Pacific, including the cooling in the Sea of Okhotsk. However, in the Indian Ocean we do not see the 1–2 8C of warming which they record following uplift. Also, there is a significant summer SST warming following uplift in the South China Sea in the study of Kitoh (2004) which does not appear to be present either in Abe et al. (2003) or our own study. All the studies agree that the change in SST is not much greater than 2 8C in any region, which is unlikely to be detectable given the current quality of SST proxies on these timescales, and the compounding factors of other possible influences on SST, such as greenhouse gas concentrations. In terms of surface air temperature, which strongly influences the vegetation distributions, our anomaly Emod –Eflat is shown in Figure 3a. The strong lapse-rate effect is clearly visible, as is an associated cooling to the south and east of the uplifted plateau, and a warming to the north. However, outside of the uplifted region the surface air temperature changes are relatively small. This is reasonably consistent with the previous study of Jiang et al. (2008), which presented a similar anomaly (their Fig. 4). They show a very similar cooling over the plateau itself of about 20 8C, with changes of the order 1–2 8C outside of the plateau. They do not have the warming to the north but instead have cooling, and a warming to the west. This is likely to be a model-dependent feature rather than a difference in experimental design-although they used fixed SSTs this is unlikely to have a major effect in the regions to the north and west of the plateau. Precipitation changes are typically found to be modeldependent (Solomon et al. 2007). Figure 3b shows the annual mean precipitation change predicted by our model following uplift, Emod– Eflat. Over the plateau, there is wettening of up to 3 mm day21. There is also a dipole of wettening and drying in the equatorial Western Pacific, perhaps associated with changes in ENSO (Kitoh 2007); although as stated in the introduction the relationship between the monsoon and ENSO is not well represented in our model. The magnitude of wettening over the plateau is consistent with the study of Jiang et al. (2008) (their fig. 6), as is the region of drying to the north. They do not predict large changes over the oceans but this is unsurprising as they used fixed SSTs. The magnitude of our precipitation changes are less than those of Kitoh (2004). This could be related to the fact that they uplifted all mountain regions, not just Tibet. In summary, our model results are broadly consistent with previous works in terms of temperature response. For precipitation, the results appear to vary more between

models, as is found for predictions of future climate change (e.g. Solomon et al. 2007). The next generation of climate models, such as those which will be used in the next assessment report of the IPCC, are more advanced than HadCM3L in several ways. Firstly, they run at higher resolution in both the atmosphere and ocean, and have a more complete representation of the Earth system, in that they include, for example, the interactive response of vegetation with climate forcing. It is possible that a higher resolution climate model would lead to a better representation of orographic effects, in particular the blocking of atmospheric flow. However, given that the orography pre-uplift is not at all well constrained, it is unlikely that such higher resolution could really bring many benefits to a palaeoclimate study, although it may improve the model’s representation of the current atmospheric state. Additionally, higher resolution models are much more computationally expensive than HadCM3L, and a balance is needed between resolution, and, given computational constraints, the tunability of the control climate and attainability of equilibrium. Resolution potentially also plays a role in the response of vegetation. As Tibet uplifted, it is very likely that there was a major change in the dominant vegetation regime, and thus changes in surface hydrology and albedo, which could have a important impact on climate evolution on a local scale. Such possibilities should be investigated in future work, although care would need to be exercised to distinguish changes in vegetation driven by global Cenozoic climatic deterioration from those caused by localized plateau surface uplift. We have made no attempt to modify the model boundary conditions to be representative of any particular period of uplift. In particular, the boundary conditions outside of Tibet are those of present day, and the vegetation is fixed. CO2 is also set at a constant 280 ppmv. In reality, the uplift of Tibet was occurring during a time when CO2 was not constant (Kurschener et al. 2008) and ice sheets were probably varying. As such, any proxy records through the time period concerned will have many influences. This study should be regarded as a sensitivity study, designed to explore the magnitude of climate change resulting from Tibetan uplift, and possible uplift histories, in isolation, rather than an attempt to simulate the uplift history of Tibet accurately. It is possible that appropriate modifications to the boundary conditions would radically alter the climatic response to the orographic forcing. However, given the similar climatic response in the model to two quite different uplift histories (Efrac and Esouth), we consider this unlikely (although potentially important on local scales).

TIBETAN UPLIFT AND PALAEOCLIMATE PROXIES

A possibly more important shortcoming in this regard is the fact that in the context of calculating runoff, we have not modified the river routing at the same time as changing the orography. The river watersheds, and indeed river mouths, may have changed dramatically over the timescales associated with uplift. Clark et al. (2004) and Brookfield (1998) are among a number of authors who have argued that uplift of eastern Tibet and SW China has caused substantial drainage re-organization in SE Asia. In particular, it has been proposed that the middle Yangtze has reversed flow, while the headwaters of the Yarlung, Salween, Mekong and Yangtze used to flow into the Red River and have subsequently been lost. Mass balance arguments concerning the amount of sediment in the Red River offshore have been used to argue in favour of major drainage re-organization (Clift et al. 2006). More recent single grain provenance work targetting grains in sandstones from the palaeo-Red River delta now rules out influence from the upper Yangtze, Mekong and Salween into the Red River (Clift et al. 2008), while leaving possible influx from the Middle Yangtze open. This revised understanding is more consistent with our model assumption. The BIOME4 model is based on the characterization of modern-day biomes, and takes no account of evolutionary changes which occurred during or since the time of uplift (e.g. Cerling et al. 1997). The plant types are sufficiently broad that it is likely that the main biome types would have existed in some form during the period of uplift, even if their bioclimatic limits may have been somewhat different. Here we are interested in the response of general vegetation classes to the uplift, rather than the exact form that the vegetation may have taken. In terms of interpreting proxies, it would be very hard to distinguish evolutionary changes which were driven directly by uplift or by other forcings.

Conclusions In this study we have carried out an ensemble of GCM experiments in which ‘snapshots’ of Tibetan uplift are simulated, from a flat plateau (Eflat) to the modern configuration (Emod) via two possible syn-uplift configurations – a wholesale uplift of the plateau (Efrac) and an uplift from the south (Esouth). We have analysed the model predictions in terms of the effect on uplift of three commonly-used palaeoclimate proxies: SST, vegetation, and river discharge. We have shown that the SST changes associated directly with uplift are very small, and are unlikely to be detectable using current methods of

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palaeo-SST reconstruction. An exception is in the subtropical Western Pacific, but the area of cooling there is still relatively small. SST shows little possibility of distinguishing between different uplift histories, as the SST differences between simulations Efrac and Esouth are small. Vegetationbased proxies unsurprisingly show a large sensitivity over the plateau itself, where temperature and precipitation changes are both relatively large. They also show potential for distinguishing the history over the plateau, as the vegetation changes over Tibet in Efrac and Esouth are very different. Outside of the plateau, the vegetation is relatively insensitive to the temperature and precipitation changes associated with uplift. There are some subtle shifts in vegetation boundaries, but if these were observed in the proxy record, it would be hard to interpret them as being directly due to uplift, as other factors could have a similar magnitude effect. River discharge also shows some sensitivity to uplift. It is important to use the most sensitive river systems, which the model indicates to be the Ganges/Brahmaputra, the Yangtze, and the Pearl. In particular, the Ganges/Brahmaputra discharge appears to be a good proxy for distinguishing different uplift histories, as the magnitude of run-off is significantly different in simulations Efrac and Esouth. This work was carried out in the framework of the British Antarctic Survey GEACEP (Greenhouse to ice-house: Evolution of the Antarctic Cryosphere And Palaeoenvironment) programme. DJL is funded by British Antarctic Survey (BAS) and Research Councils UK (RCUK) fellowships.

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Climate modelling study on mountain uplift and Asian monsoon evolution A. KITOH*, T. MOTOI & O. ARAKAWA Meteorological Research Institute, 1-1 Nagamine, Tsukuba, Ibaraki 305-0052, Japan *Corresponding author (e-mail: [email protected]) Abstract: Impacts of mountain uplift on the Asian monsoon and adjacent seas are investigated by climate model sensitivity studies. Two sets of general circulation model (GCM) experiments are performed. Using an atmosphere-ocean coupled GCM, a progressive mountain uplift experiment is performed. During boreal summer, monsoon precipitation is confined in the deep tropics around 108N in the no-mountain case, but as mountains become higher, heavy rain areas move inland from the East Asian coast with stronger upward winds and increased rainfall over the southeastern Tibetan Plateau region. An increase of freshwater discharge from the Asian rivers results in a significant decrease of sea surface salinities over the Bay of Bengal, the South China, East China and Yellow Seas. A high-resolution atmospheric GCM experiment, which shows improvement in reproducing the present-day model climatology, gives more precise information on precipitation and the circulation changes caused by mountain uplift.

Ocean and land-surface features in the climate system have been altered by tectonic events through the geological history of the Earth. Tectonic evolution of the land-sea distributions, ocean gateways and mountain uplift are considered to be key factors of global climate changes together with the Milankovitch solar forcing and greenhouse-gas variations (Crowley & North 1991; Ruddiman 1997; Crowley & Burke 1998; Zachos 2001; Clift & Plumb 2008). By comparing the results of numerical experiments with and without the effects of mountains in an atmospheric general circulation model (AGCM), Manabe & Terpstra (1974) found that the mountains affect hydrological processes by modifying the field of threedimensional advection of moisture, and significantly alter the global distribution of precipitation. Hahn & Manabe (1975) also found that the presence of the Tibetan Plateau in the AGCM was required for northward extension of a monsoon climate onto the Asian continent. These findings suggest the possible linkage between mountain uplift and monsoon evolution. Ruddiman et al. (1989) assembled geological evidence for mountain uplift and climate change to compare them with the results from AGCM sensitivity experiments with and without mountains by Kutzbach et al. (1989). According to Ruddiman & Kutzbach (1989) the comparison indicated qualitative consistency between geological evidence and AGCM estimation, concerning the influence of mountain uplift on the monsoon system. Palaeoclimate implications for the possible linkage between mountain uplift and monsoon evolution

were also discussed based on AGCM sensitivity experiments (Manabe & Broccoli 1990; Broccoli & Manabe 1992; Liu & Yin 2002). The results from Manabe & Broccoli (1990) and Broccoli & Manabe (1992) implied that expansion of midlatitude arid regions during the late Cenozoic might be associated with monsoon evolution induced by mountain uplift. Liu & Yin (2002) pointed out that the evolution of the East Asian monsoon might be more sensitive to the uplift of the Tibetan Plateau than that of the South Asia monsoon. By using atmosphere-mixed layer ocean coupled models (AGCM-MLOs), the effects of the Himalayan and Tibetan Plateau uplift on the Asian monsoon were investigated (Prell & Kutzbach 1992; Kutzbach et al. 1993; Kitoh 1997). Comparing the AGCM-MLO results with palaeoclimate and palaeocean reconstructions, Prell & Kutzbach (1992) indicated that Asian monsoon precipitation increased as the plateau elevation was increased and hence runoff increased. Kutzbach et al. (1993) and Kitoh (1997) revealed the fall of sea surface temperatures (SSTs) in the western North Pacific marginal seas due to mountain uplift. Using idealized stepwise increase of plateau elevation in a climate model, An et al. (2001) obtained results supporting the argument that the stages of evolution of Asian monsoons were linked to the phases of Himalaya-Tibetan plateau uplift and to Northern Hemisphere glaciations. Recently ocean response to mountain uplift was studied by using atmosphere-ocean-land surface coupled models (AOGCMs) (Kitoh 2002, 2004,

From: Clift, P. D., Tada, R. & Zheng, H. (eds) Monsoon Evolution and Tectonics –Climate Linkage in Asia. Geological Society, London, Special Publications, 342, 293–301. DOI: 10.1144/SP342.17 0305-8719/10/$15.00 # The Geological Society of London 2010.

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2007; Abe et al. 2003, 2004, 2005). In particular, Kitoh (2004) showed that the climate system response is modulated by air–sea interaction, resulting in a stronger sensitivity in monsoon rainfall in a coupled climate system than in the atmosphere-alone experiment. Kitoh (2004) also found that there are significant changes in the East Asian climate when the mountain reached about 60% of the present elevation. Jiang et al. (2008) confirmed the 60% criteria by a different climate model. Modulations of El Nin˜o/Southern Oscillation (ENSO) by mountain uplift are investigated by Kitoh (2007), where model ENSO becomes systematically weaker, shorter and less periodic when the mountain height increases. Although the AOGCM was used in the previous studies, few analyses have been carried out on the ocean’s response due to mountain uplift except for the SST changes and ENSO modulation. Obviously further analyses should be completed because mountain uplift influences not only atmospheric but also oceanic circulation. The horizontal resolution of the model is also influential. A typical horizontal resolution of AGCMs used in the above-cited studies is about 280 km. There are studies showing that the horizontal resolution improves some features of the present-day Asian monsoon climate simulation (Kobayashi & Sugi 2004; Kitoh & Kusunoki 2008). Therefore, experiments using higher resolution models are needed to study impacts of mountain uplift on monsoon climate further. In this paper, we compare the effect of model’s horizontal resolution using two sets of GCM experiments. One is a progressive mountain uplift experiment using a low-resolution AOGCM, the other is a mountain, half-mountain and no-mountain experiment using a highresolution AGCM. The former has a horizontal resolution of 280 km, whereas the latter has a resolution of 120 km. This opens the door for higher resolution modelling in palaeoclimate studies.

Model and experimental method For a progressive mountain uplift experiment, we used the global ocean-atmosphere coupled GCM (MRI-CGCM2, Yukimoto et al. 2001). The horizontal resolution in the atmosphere is about 280 km in longitude and latitude. The vertical configuration consists of 30-layer sigma-pressure hybrid coordinate with the top at 0.4 hPa. The ocean general circulation model (OGCM) has a global domain with the horizontal grid spacing of 2.58 in longitude and 2.08 in latitude with a finer meridional scale in the tropics. There are 23 vertical levels with the bottom at 5000 m. The uppermost layer has a 5.2 m thickness. The atmosphere and

the ocean interact with each other by exchanging fluxes of heat, freshwater and momentum at the sea surface. The fluxes are exchanged every 24 hours in the model. The control run (M100) was integrated for 50 years with a realistic land–sea distribution and mountain height. The elevation of the highest grid point over the Tibetan plateau is 5536 m in M100. The model climatology of this control run is described by Kitoh (2004). In the M000 run, the worldwide mountain height was set to zero, but keeping the land –sea distribution the same. The M020, M040, M060 and M080 runs used the 20%, 40%, 60% and 80% of the M100 mountain height, respectively. We also included the M120 and M140 runs that used the increased mountain height of 120% and 140%. All runs were performed for 50 years integration and the last 40 years data from these eight runs are analysed. In the second experiment, we used a highresolution AGCM, where the horizontal resolution is about 120 km. Three runs were performed with full, half and no mountains; they are denoted as A100, A050 and A000, respectively. The AGCM is integrated for 12 model years each, and the last 10 years mean are used for analysis.

Atmosphere-ocean coupled modelling results Figure 1 shows the geographical distributions of December-January-February (DJF) and June-JulyAugust (JJA) mean surface winds in the M020, M060 and M100 together with the present-day observed climatology. Observed winter (DJF) surface winds are characterized by northwesterly winds around Japan and northeasterly winds over SE Asia and South Asia. All the experiments show common features with westerly winds north of 308N and easterly winds south of 308N, but their intensity varies between the experiments. Northeasterly winds over the Arabian Sea and the Bay of Bengal are rather strong in lower mountain runs, and become weaker with mountain uplift. Around Japan, the surface wind blows from west in M020, but it blows from NW in M060 and M100 with weaker intensity. Thus, the Asian winter monsoon wind becomes weaker with mountain uplift, which is in opposite with the summer monsoon case. In JJA, the observation shows a strong southerly cross-equatorial wind along the eastern coast of Africa, which becomes the Somali jet reaching the Arabian Sea and India. There also is a wind maximum over the Bay of Bengal. Over the Pacific, the trade winds (easterly winds) dominate, converging with the monsoon westerlies around

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Fig. 1. Geographical distributions of December– January–February (DJF) mean and June–July–August (JJA) mean surface winds (m s21). (a) Observation DJF; (b) 20% DJF; (c) 60% DJF; (d) 100% DJF; (e) Observation JJA; (f ) 20% JJA; (g) 60% JJA; and (h) 100% JJA.

the Philippines, and then flow toward southern China and Japan. The model (M100) reproduces well the observed Asian summer monsoon wind system. These features are generally seen in all the mountain runs except with different distributions and intensity. In M000, westerly winds over the Eurasian continent and southwesterly winds over the ocean off the east of Japan are strong (not shown). These winds become weak in M020 and M040. Over the Indian Ocean, the mountain uplift intensifies the Somali jet, the westerly winds in the

southern Bay of Bengal and also the along-shore winds off the Sumatra-Java islands. Over the Arabian Sea, the southerly winds generally dominate. However, in the runs with lower mountains (M000, M020 and M040), northerly winds dominate even in summer in the northern part of the Arabian Sea. Figure 2 shows the geographical distributions of DJF and JJA mean precipitation in the M020, M060 and M100 together with the present-day observed climatology. In winter, mountain uplift has a large

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Fig. 2. Geographical distributions of December– January–February (DJF) mean and June– July– August (JJA) mean precipitation (mm d21). (a) Observation DJF; (b) 20% DJF; (c) 60% DJF; (d) 100% DJF; (e) Observation JJA; (f) 20% JJA; (g) 60% JJA; and (h) 100% JJA.

effect on the precipitation change over the region from southern China to the off the east coast of Japan, where the precipitation due to synoptic disturbances becomes intensified. The location of large precipitation moves inland in China with mountain uplift. Precipitation over the western Indian Ocean decreases with mountain height associated with weakening of northeasterly winter monsoon winds. The summertime precipitation drastically changes with mountain uplift. In the no-mountain (M000) run, a heavy precipitation belt is confined

in the deep tropics within 108 latitude. To the north, a dry region, with precipitation less than 1 mm d21, almost entirely covers the Eurasian continent in the 258N – 408N zone, yielding a dry climate even in East Asia. In M020, M040 and M060, the precipitation increases over land in SE Asia and East Asia, and a dry area retreats westward. In M060, a rain band corresponding to the Baiu is seen from Taiwan to southern Japan. A northward shift of large precipitation area continues into M080 and M100. The dry area over the western Arabian Sea becomes distinct with

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Fig. 3. June–July–August mean precipitation (mm d21) as a function of mountain height (%) for the South Asian Monsoon region (SAM: 708E –1008E, 108N–308N: thick solid), the Southeast Asian Monsoon region (SEAM: 1008E– 1308E, 58N–258N: thick dashed), and the East Asian Monsoon region (EAM: 1208E –1408E, 258N –358N: thin solid).

mountain uplift, and thus an east –west precipitation contrast appears in the Indian Ocean in moderate to high mountain cases. Figure 3 shows the sensitivity of the JJA mean precipitation as a function of mountain height in the three monsoon regions, that is, the South Asian Monsoon region (SAM: 708E –1008E, 108N– 308N), the SE Asian Monsoon region (SEAM: 1008E–1308E, 58N–258N), and the East Asian

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Monsoon region (EAM: 1208E– 1408E, 258N – 358N). The SAM precipitation amount almost increases linearly with mountain height. The EAM precipitation plateaus off at M100 and M120, and shows a slight decrease in M140. The SEAM precipitation has the maximum at M080, and further mountain uplift results in a decrease in precipitation amount because intense rainfall area moves inland. Up to 80% mountain uplift, all the Asian summer monsoon region would experiences increase in rainfall with mountain uplift, implying that freshwater discharge from the Ganges, the Red, Mekong, Yellow and Yangtze rivers increases with progressive mountain uplift due to the movement of intensified rainfall area into their drainages. Ocean surface conditions (SST and sea surface salinity) as well as ocean general circulation and its interannual variability also should change with mountain uplift (Kitoh 2004, 2007), because the mountain-induced atmospheric circulations affect heat, momentum and water fluxes at the sea surface. Figure 4 shows the DJF and JJA mean SST distributions in M020 and M100. The M100 case reproduces the observed seasonal variation except for excessive intrusion of cold water at the equatorial Pacific into the western Pacific. The characteristic feature in the differences between M100 and M020 is that the M100 SST is warmer than in M020 over the central/eastern Indian Ocean and the western tropical Pacific, whereas the M100 SST is colder over the western Indian Ocean. The former is associated with changes in

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Fig. 4. Geographical distributions of December– January–February (DJF) mean and June–July–August (JJA) mean sea surface temperature (8C). (a) 20% DJF; (b) 100% DJF; (c) 20% JJA; and (d) 100% JJA.

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the intensity of the subtropical anticyclone and trade winds (Kitoh 2004). The latter is associated with more evaporation and larger upwelling of deep cold water in the western Arabian Sea due to stronger winds. The lower mountain cases have weak Somali jet, and thus the SST does not become lower. The SST distribution is zonally uniform in the M020, whereas it has a large zonal gradient in M100 (and in the observations). Coastal upwelling off Somalia and downwelling off Sumatra-Java are enhanced in summer due to mountain uplift, resulting in a dipole sea surface water change across the tropical Indian Ocean, colder in the west and warmer in the east. Larger offshore and onshore Ekman transport induced by stronger alongshore winds off Somalia and off Sumatra-Java during the summer monsoon in 100% experiment cause enhanced upwelling and downwelling in these regions, respectively (see Fig. 1). Changes in precipitation intensity and its distribution bring about changes in sea surface salinity (SSS). Figure 5 shows the DJF and JJA mean SSS distributions in M020 and M100. In M100 (and in the observations) there is low salinity water around the maritime continent and high salinity water in the Arabian Sea. The SSS in the Bay of Bengal is lower than that in the Arabian Sea. The low SSS in the Yellow Sea and the East China Sea is also noted in M100. Such a spatial feature is not seen in M020, where the SSS distribution is spatially uniform. In particular, there is almost no contrast in SSS between the

Arabian Sea and the Bay of Bengal. An increase in freshwater discharge from the Ganges, the Red, Mekong, Yellow and Yangtze rivers results in a significant decrease of SSS in the Bay of Bengal, the South China, East China and Yellow Seas throughout the year. The SSS in the Sea of Japan also decreases in the southern region due to increase of freshwater transport by Tsushima current through the Tsushima Strait. The SSS in the Arabian Sea increases due to decrease of in situ precipitation. In summary, we have shown that systematic changes occurred in precipitation and circulation fields with progressive mountain uplift: the precipitation area moved inland with mountain uplift, whereas the Pacific subtropical anticyclone and associated trade winds became stronger with mountain height. The mountain uplift also resulted in SST and SSS changes. About half of the mountain height (60% height of the present in our experiment) is found to be needed to make the Asian monsoon climate similar to the present conditions. The Asian monsoon does exist with lower mountain cases; the regional distribution and intensity of the monsoon ingredients (i.e. precipitation and wind system and so on) depend on the mountain height.

High-resolution atmospheric modelling results The experiment described in the former subsection has been performed with an AOGCM, but its

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Fig. 5. Geographical distributions of December– January–February (DJF) mean and June– July– August (JJA) mean sea surface salinity (psu). (a) 20% DJF; (b) 100% DJF; (c) 20% JJA; and (d) 100% JJA.

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horizontal resolution is rather coarse with about 280 km and is not high enough to resolve the realistic land-sea distribution, topography and steep mountains. Those coarse resolution models only resolve a bowl-like Tibetan Plateau, but have no Himalayan Ranges nor Burmese Mountains, to name a few. Because the geographical distribution of precipitation depends largely on any given model’s horizontal resolution, it is desirable to use the highest resolution possible. Here we use a highresolution AGCM with a 120 km grid size to investigate mountain uplift effects on the Asian monsoon. Results with mountain (A100), half-mountain (A050) and no-mountain (A000) runs are shown. Figure 6 shows the geographical distributions of DJF and JJA mean precipitation in the A000, A050 and A100. Comparing Figure 6c and 6f with Figure 2d and 2h, one can see an improvement of present-day precipitation climatology against the observations (Fig. 2a, e). For example, a lowresolution AOGCM has an excessive precipitation

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at the southeastern corner of the Tibetan Plateau in JJA, but this is remedied by a high-resolution AGCM. A windward precipitation maximum and a leeward dry area are simulated well around the southern part of India (Fig. 6f). A summertime rain belt at the foothill of the Himalayas is seen in the high-resolution AGCM. It falls as snow in winter, which is clearly seen in Figure 6c. During boreal winter, effects of mountains on precipitation appear distinctly in the high-resolution model experiment. Without mountain, synoptic disturbances easily move into the Eurasian continent from the Atlantic but their activity becomes weak and precipitation is reduced over the continent (Broccoli & Manabe 1992). This is seen both in the coarse-resolution model (Fig. 2) and in the high-resolution model result (Fig. 6). In our highresolution case, mountain uplift makes cyclonic activity more active with more precipitation in the Mediterranean region in DJF. Increased precipitation is simulated over mountainous regions over

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Fig. 6. Geographical distributions of December– January–February (DJF) mean and June–July–August (JJA) mean precipitation (mm d21) by the 120-km mesh atmospheric GCM. (a) 0% DJF; (b) 50% DJF; (c) 100% DJF; (d) 0% JJA; (e) 50% JJA; and (f ) 100% JJA.

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Turkey and around. A crescent-shaped precipitation over the Zagros Mountains appears as mountains become higher. The reproducibility of this precipitation distribution is important in future climate projections to assess future water resources (Kitoh et al. 2008). The same is true in the Himalayas where a precipitation belt is clear in A100, is barely seen in A050, but does not exist in A000. Peninsular India obtains about 1 mm d21 precipitation in A100 and A050, but becomes dry in A000. Activities of synoptic disturbances affect precipitation over the Yangtze River basin in southern China, where ample winter precipitation only appears in the full-mountain case. During the summer monsoon season, as mountains become higher, heavy rain bands become sharper and move into the East Asian land region as in the coarse-resolution experiment. However, in this high-resolution experiment, a reduction in precipitation over the Tarim Basin is clearly seen as the summer precipitation over Tian Shan Mountain appears in A050 and becomes distinct in A100. Another difference between the two experiments with different resolutions is that there is some amount of rainfall over southern China even in the no-mountain run of the high-resolution experiment. A part of such a difference should be associated with different SST between the two experiments whether the ocean is interactively coupled or not. Kitoh (2004) showed that the no-mountain AGCM had more precipitation over SE Asia than the no-mountain AOGCM.

induce a change in the global ocean thermohaline circulation with resultant drastic changes in sea-ice distribution in the polar oceans. The ocean stratification becomes stronger due to the freshening in the marginal seas. The changes in the ocean stratification and vertical velocity (upwelling and downwelling) lead the variations of the nutrients in the ocean. It strongly influences the marine ecosystem including the phytoplankton, which is a good proxy for palaeoceanographic reconstruction. Land vegetation and oceanic plankton may be influenced by mountain uplift. Sensitivity experiments by Yasunari et al. (2006) revealed that the role of land surface processes might be similar to those of large-scale mountains on the monsoon climate. The linkage between the results from the International Continental Scientific Drilling Program and Integrated Ocean Drilling Program is very informative for the model studies. In the high-resolution model, we found a large spatial contrast in precipitation changes by mountain uplift around the coastal regions. When the high-resolution AOGCM is used, such a precipitation change in addition to more realistic river flow would lead to different sea surface salinity changes over the coastal seas in the South China Sea and the Arabian Sea. Because the sign of the precipitation change can be reversed simply by which model resolution was used, caution is needed to make a comparison between simulation results and palaeoclimate proxies. Therefore, highresolution AOGCM experiments are desirable for further studies.

Discussion and conclusions

This work was carried out under ‘Modeling Study of Climate System and its Variability’ at MRI, and was also supported by the Grant-in-Aid for Scientific Research (C20540433).

In this study, impacts of mountain uplift on the Asian monsoon are investigated with two different horizontal resolution models. Finer structure due to highly resolved topography in 120 km mesh is clearly established in the precipitation distributions as mountain height increases. There is a large change in the distribution of dry climates (desert and steppe). Area coverage of dry climates is the largest in the no-mountain case and decreases in its area extent with mountain uplift. The finer structure of precipitation in the atmospheric GCM with higher resolution also has a finer convergence and divergence structure in the sea surface wind fields. Wind stress must influence ocean current and then SST, especially in the frontal region around the western boundary current such as the Kuroshio and the Oyashio in the North Pacific and the Somali current in the Indian Ocean. The heat transported by ocean currents is comparable in size to that carried by the atmosphere. The temperature and salinity changes caused by changes in the wind-driven ocean circulation may

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MODELLING OF MOUNTAIN UPLIFT AND MONSOON Broccoli, A. J. & Manabe, S. 1992. The effects of orography on middle-latitude northern hemisphere dry climates. Journal of Climate, 5, 1181– 1201. Clift, P. D. & Plumb, R. A. 2008. The Asian Monsoon: Causes, History and Effects. Cambridge University Press, Cambridge. Crowley, T. J. & Burke, K. C. (ed.) 1998. Tectonic Boundary Conditions for Climate Reconstruction. Oxford University Press, New York. Crowley, T. J. & North, G. R. 1991. Paleoclimatology. Oxford Monographs on Geology and Geophysics, 18. Hahn, D. G. & Manabe, S. 1975. The role of mountains in the south Asian monsoon circulation. Journal of the Atmospheric Sciences, 32, 1515–1541. Jiang, D., Ding, Z., Drange, H. & Gao, Y. 2008. Sensitivity of East Asian climate to the progressive uplift and expansion of the Tibetan Plateau under the Mid-Pliocene boundary conditions. Advances in Atmospheric Sciences, 25, 709–722. Kitoh, A. 1997. Mountain uplift and surface temperature changes. Geophysical Research Letters, 24, 185– 188. Kitoh, A. 2002. Effect of large-scale mountains on surface climate—a coupled ocean-atmosphere general circulation model study. Journal of the Meteorological Society of Japan, 80, 1165– 1181. Kitoh, A. 2004. Effect of mountain uplift on East Asian summer climate investigated by a coupled atmosphere-ocean GCM. Journal of Climate, 17, 783–802. Kitoh, A. 2007. ENSO modulation by mountain uplift. Climate Dynamics, 28, 781–796. Kitoh, A. & Kusunoki, S. 2008. East Asian summer monsoon simulation by a 20-km mesh AGCM. Climate Dynamics, 31, 389–401. Kitoh, A., Yatagai, A. & Alpert, P. 2008. First super-high-resolution model projection that the ancient “Fertile Crescent” will disappear in this century. Hydrological Research Letters, 2, 1– 4. Kobayashi, C. & Sugi, M. 2004. Impact of horizontal resolution on the simulation of the Asian summer monsoon and tropical cyclones in the JMA global model. Climate Dynamics, 23, 165 –176. Kutzbach, J. E., Guetter, P. J., Ruddiman, W. F. & Prell, W. L. 1989. Sensitivity of climate to late

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Index Page numbers in italic denote figures. Page numbers in bold denote tables. Aertashi section 68 Cenozoic sediments 69, 73 facies associations 69, 71, 72, 74– 76 lithostratigraphy 69, 71 magnetostratigraphy 69, 71 fan delta 72, 73 Ailao Shan-Red River Fault Zone 221, 223, 238 Aka Aiteputh Formation 110, 112 climate change 120–122 depositional facies 110– 112, 116 stratigraphy 113, 114 Alaknanda-Bhagirathi River 131, 133 CO2 consumption 144, 147 silicate weathering 132, 133, 139, 144, 146 strontium isotopes 135, 141, 142, 143, 145 alkenones 9, 10 Altyn Tagh Fault 62 Antarctica, glaciation, and central Asian aridification 34–35 apatite, fission track analysis 191, 196, 203, 204, 205 40 Ar/39Ar mica dating, Indus basin sediments 191, 197– 198, 203– 205, 206, 207 Arabian Sea monsoonal upwelling 153–154 ODP Hole 716A 18, 19, 24 aridification central Asia 29– 41, 45, 257 stepwise drying 31– 35 see also desertification Artux Formation 46, 48– 50, 52, 54, 69 age model 52, 53, 54 grain size analysis 57– 59 linear sedimentation rate 54, 56, 60– 61 lithostratigraphy 70, 71, 72, 73, 75 provenance 59, 60 Asia, central, aridification 29–41, 257 Asian monsoon evolution, and mountain uplift, modelling 293 –300 Atlantic Multidecadal Oscillation 18 Atlantic Ocean and Central American Seaway 267, 272 and Indian monsoon 18 thermohaline circulation 18 Bap-Malar playa 169–170, 170 Bashiblake Formation 71, 72 Beiguoyuan sections, OSL dating 91– 96, 100–102 Belan River, Quaternary fluvial record 163, 164, 165, 176 Bengal Basin, Holocene monsoon intensity 177 Bengal Fan organic carbon 202, 203 sedimentation 257 Bering Sea, sea ice 267, 269– 270, 272, 276 BIOME4 model 281, 283, 285, 289 Black Loam Formation 91, 97, 98

Brahmaputra River 131, 133 CO2 consumption 144, 147 run-off 286, 289 silicate weathering 132, 133–135, 136, 139, 144, 146 strontium isotopes 135, 141, 142, 145 CaCO3 precipitation 129 –130 variation, summer monsoon, Chinese Loess Plateau 88, 94– 102, 103, 104, 105 calcrete Aka Aiteputh Formation 110, 112, 116, 121 Thar Desert and Ganga River plains 167–168, 176 carbon, organic Bengal Fan 202, 203 Indus basin sediments 191, 195, 199–202, 213 Cenozoic Late, aridification, central Asia 29–41 Tarim Basin sediments 68–77 Central American Seaway 265 –276 closure, palaeoclimatic implications 272, 274 and East Asian monsoon 265, 271 effect on thermohaline circulation 267 Chang Jiang River 131, 133 CO2 consumption 147 run-off 286, 287, 289 silicate weathering 132, 133, 134, 136, 144, 146 strontium isotopes 141, 142, 145 chert, Aka Aiteputh Formation 110 –111, 116 Chinese Loess Plateau 8, 30, 46 Aeolian sediments 32, 33–34, 89 age models 88, 90 East Asian summer monsoon 87–106, 256–257 OSL dating 90–106 grain size, and orbital variation 80– 82, 84 sedimentation rate 104–105 circulation, thermohaline North Atlantic 18 North Pacific 267, 269– 271, 275 –276 climate modelling Central American Seaway 266– 276 mountain uplift 293– 300 CO2 atmospheric, silicate weathering 129– 149 Himalaya-Tibetan Plateau consumption rates 134– 135, 143– 144 sink potential 130– 149 conglomerate Samburu Hills 110, 114, 117 Tarim Basin 48– 52, 54, 69, 72, 75– 76, 77 cooling, central Asia 29, 34– 41 Cymbaloporetta squammosa 20–25 Dehra Dun Re-entrant 156, 162 delta deposits, Samburu Hills 117 desertification, Tarim Basin 46–63

304 Didwana playa 169, 170, 171 dunes, aeolian, Tarim Basin 72, 74–75, 76– 77 East Africa, Miocene 109– 124 East Asian monsoon and Central American Seaway 265– 276 evolution 7– 10 and Himalaya–Tibetan Plateau uplift 45 intensification, and weathering 238, 239–240, 255, 256 reconstructions 221– 222, 233, 237, 240, 256 summer 8– 9, 87– 106 Central American Seaway 271, 274, 276 evolution 256–257, 258 global ice volume 102– 104 OSL dating 90–91 Quaternary, radiometric dating 90–106 South China Sea 245–260 tectonics, and continental erosion 219– 240 variability 10– 11 marine records 5– 14 winter Central American Seaway 271, 274, 276 evolution 258 and global ice volume change 79– 84 influence of larger sea ice 269–271 South China Sea 10, 11 East China Sea East Asian summer monsoon 11 productivity 11, 13 sediment cores 7, 11 El Nin˜o-Southern Oscillation and Indian monsoon 18 in modelling 280–281, 288 elemental analysis, Indus basin sediments 191, 193– 195 erosion, and monsoons 1, 10 East Asian monsoon 219–240 Western Himalaya 210–211 fan delta, Tarim Basin 72, 73 fan-terrace systems, Quaternary, India 161–162 fans, alluvial, Tarim Basin 47, 61– 62 fission track analysis, Indus basin sediments 191, 196, 203, 204, 205 foraminifera ODP Hole 716A 20–25 South China Sea 10, 12, 13 Gandak River 131, 133 CO2 consumption 147 silicate weathering 132, 133, 136, 139, 141, 144, 146 strontium isotopes 141, 142, 143, 145 Ganga River 131, 133 plains calcrete 168, 176 fluvial records 162–165, 174–175, 176 monsoon intensity 177 Quaternary lacustrine records 171 run-off 286, 287 silicate weathering 132, 133, 134, 136, 139, 141, 144, 146 strontium isotopes 135, 141, 142

INDEX Ghaghara River 131, 133 CO2 consumption 144, 147 silicate weathering 132, 133, 135, 136, 139, 141, 144, 146 strontium isotopes 135, 141, 142, 143, 145 glaciation Antarctica 34–35 and central Asian aridification 34–41 northern hemisphere 8, 10, 35, 80, 84 and productivity, South China Sea 10 Globigerinoides ruber 20, 21, 22–23, 24, 25, 221 Globigerinoides sacculifer 20, 21, 22, 24, 25, 221 Gobi, Aeolian sediments 29, 31–32, 76 Gularchy sediments 188 strontium isotope analysis 188, 192 HadCM3L model 280 –281 Hainan Island, erosion 220, 223, 231, 239 halocline North Pacific 267, 269– 271, 272 sea ice formation 265 Haripur Kohl, palaeosols 158, 173 Himalaya Greater tectonics and the monsoon 211, 213 zircon dating 205– 207, 208 uplift 46 western, monsoon and erosion 210–211 Himalaya–Tibetan Plateau 131 CO2 sink potential 130– 149 CO2 uptake chemical weathering rates 140–141, 143 consumption rate 143–144 rivers, strontium isotope ratios 135, 138, 141, 142, 143 uplift and Asian monsoon 7, 45, 68, 153, 173 timing 46 see also Tibetan Plateau Holocene monsoon intensification 162, 165, 166, 167, 171– 172, 175, 177, 257 Indus River basin 186–214 OSL dating 91, 92, 93, 94– 97, 99, 100, 106 ‘Holocene Optimum’ 90, 100, 106 Hong River 131, 133 CO2 consumption 147 silicate weathering 132, 133, 134, 136, 144, 146 strontium isotopes 141, 142, 145 Huang He River 131, 133 CO2 consumption 147 run-off 286, 287 silicate weathering 132, 133, 134, 136, 144, 146 strontium isotopes 141, 142, 145 ice volume, global East Asian summer monsoon 102–104 East Asian winter monsoon 79–84 IMAGES cores MD012404 11, 13 MD972142 10, 12 India fan-terrace systems 161– 162 western, Quaternary fluvial records 165– 167

INDEX Indian monsoon 17, 18 summer 17–18, 154 evolution 153–177, 257, 258 intensification 162, 165, 166, 167, 171–172, 175, 177, 257 Kenya Rift 109, 122– 124 Late Quaternary records 161– 171 pollen data 168, 172 –173 winter 17, 18 Indian Ocean climate variability 17– 25 monsoons 17, 19 productivity 257 Indonesian seaway, closure 8, 9 –10 Indus River 131, 133 basin 40 Ar/39Ar mica dating 191, 197– 198, 203– 205, 206, 207 Asian monsoon intensification 186– 214 environment 213 fission track analysis 191, 196, 203, 204, 205 mineralogy 187–188, 198–199, 200 organic carbon 191, 195, 199– 202, 213 sediment analysis 186– 209 sediment variability 212, 257 strontium isotopes 188, 191, 192, 199, 201 U– Pb zircon dating 191, 198, 205– 207, 208, 209 CO2 consumption 144, 147 silicate weathering 132, 133, 134, 136, 139, 144, 146 strontium isotopes 135, 141, 142, 145 insolation, and East Asian Monsoon 11, 13, 90, 102 Intertropical Convergence Zone 18 Kenya Rift climate 122–124 Jati sediments 188 U–Pb zircon dating 198, 208, 209 Kalpi carbonates 168 Quaternary fluvial record 163, 164, 165 Kangra Sub-basin, palaeosols 158, 159 Keliyang section 68 Kenya Rift 110, 111 climate change 120–124 Indian summer monsoon 122– 124 Keti Bandar 186, 189, 190 Kongia Formation 110, 112, 113 Kosi River 131, 133 CO2 consumption 144, 147 silicate weathering 132, 133, 135, 136, 139, 144, 146 strontium isotopes 135, 141, 142, 143, 145 Kunlun Mountains 47, 68 alluvial fans 62 loess deposits 46, 47 yellow siltstone 48 Kuroshio Current 265, 267, 272, 275– 276 La Nin˜a conditions 9– 10 Last Glacial Maximum Ganga River basin 165, 174 Indus River basin 186, 199, 202, 203, 205, 212 Thar Desert 167 Linxia fluvial-lacustrine deposit 32

305

loess China aridification 29, 31–34, 35 magnetic susceptibility 79–80 see also Chinese Loess Plateau mountain 46, 47, 59 Luni River, Quaternary 165, 166, 176 Lunkaransar playa 169, 170, 171 Luonan, aeolian deposits 39 magnetic susceptibility analysis, Chinese Loess Platform 79, 88, 91, 94–100, 101– 105, 106 magnetostratigraphy Siwalik Group 156, 157 Tarim Basin 69, 70 Mahi River, Quaternary fluvial records 165–166, 176 Maldives Ridge, ODP Site 716A 19 equatorial wind intensity 25 proxy data 18, 20–25 Mekong River 131, 133 CO2 consumption 147 run-off 286, 287 silicate weathering 132, 133, 134, 136, 144, 146 strontium isotopes 141, 142, 145 mica, 40Ar/39Ar dating, Indus basin 191, 197–198, 203–205, 206, 207 Mid-Brunhes Climatic Event 21, 25 Miocene Aertashi Section 72, 74–75 East Africa 109 –124 East Asian summer monsoon 256–257, 259 modelling climate, Central American Seaway 266– 276 mountain uplift, and Asian monsoon evolution 293 –300 Tibetan uplift 279– 289 MOSES-2.1 model 280 mountain uplift, and Asian monsoon evolution, modelling 293– 300 MRI-CGCM2 model 294 mudstone Samburu Hills 110, 117, 118, 119 Tarim Basin 48, 69 Nachola Formation 110, 112, 113 Nakali, climate 122, 123 Namurungule Formation 110, 112 climate change 120– 122 depositional facies 115, 117, 118 stratigraphy 113, 114 Narmada river, Quaternary fluvial records 166 Nd isotope data, Indus basin 188, 192, 201, 210– 211 Neogene, Tarim Basin 48– 52, 69, 72 Neogloboquadrina dutertrei 20, 21, 22, 24, 25, 221 Ngorora Formation 121 North Equatorial Current 267, 275–276 Ocean Drilling Program (ODP) Sites 716A 18, 19, 19, 20– 25, 24 722 257 1143 7, 8 1146 221, 234, 236, 237, 246– 260 1148 7, 221–222, 236, 237, 252, 254 Okhotsk Sea, sea ice 267, 269, 270, 272, 276

306

INDEX

Okinawa Trough, sediment cores 11, 13 Oligocene Aertashi Section 72 Antarctic glaciation 34–35 optically stimulated luminescence (OSL) dating Chinese Loess Plateau 90–106 oxygen isotope data ice volume change, and orbital variation 80–82, 84 Indian summer monsoon, Siwalik Group 159– 161, 171–172, 173–174 oxygen minimum zone 18, 20 Pacific Ocean North, and Central American Seaway 265, 267–276 Subarctic halocline 267, 269 –271, 272 sea ice 267, 269–271 western sea surface salinity and temperature 5, 6 palaeosols Chinese Loess Plateau 32 Siwalik Group 158–159, 173 Panama isthmus, closure 8, 9– 10 Pearl River, run-off 286, 287 Pearl River Basin 220, 222, 231, 254 weathering records 233–234, 237 pedogenesis, summer monsoon, Chinese Loess Plateau 88, 98 playa deposits Samburu Hills 117, 119 Tarim Basin 74, 76 Thar Desert 168 –171 Plio-Pleistocene Aertashi Section 75–76 East Asian winter monsoon 80– 84 forcing mechanisms 83– 84 pollen, Indian summer monsoon 168, 172– 173 Potwar Plateau 156, 158 oxygen isotope data 159, 160 precession cycles, East Asian Monsoon 11, 13 precipitation Chinese Loess Plateau 88, 90, 100 Kenya Rift 121, 122– 123 mountain uplift model 295– 297, 299– 300 Pacific, Central American Seaway 271, 274, 275 Tibetan Plateau uplift modelling 284, 286, 288 productivity East China Sea 11, 13 Indian Ocean 257 Maldives Ridge 24, 25, 257 South China Sea 9, 10–11, 13 pteropods, ODP Hole 716A 20– 25 Qin’an loess deposit 32, 33, 36 Qiongdongnan Basin 220–221 sedimentation 231, 232, 233, 234, 235, 238, 239 seismic profiles 222, 228 structure and tectonics 228, 230–231 Quaternary aeolian records, Thar Desert 167 fan-terrace systems, India 161 –162 fluvial records Ganga plains 162 –164 western India 165–167

lacustrine records, Thar Desert and Ganga plains 168– 171 Red Clay, China 29, 31–34, 35 Red River drainage reorganization 220, 289 sediment flux 220, 222, 233, 238, 239, 240 Red River Fault Zone 221, 223, 238 Red River Shear Zone 239 red soil beds, Aka Aiteputh Formation 110, 112, 116 run-off, Tibetan Plateau uplift modelling 286, 287, 289 Sabarmati River, Quaternary fluvial records 165 –166, 176 salinity sea surface mountain uplift model 298 North Pacific, and Central American Seaway 267, 272, 275 South China Sea 8 western Pacific 5, 6 subsurface distribution, North Pacific 267 Salween River 131, 133 CO2 consumption 147 run-off 286, 287 silicate weathering 132, 133, 134, 136, 144, 146 strontium isotopes 141, 142, 145 Sambhar playa 168, 170, 171 Samburu Hills, Kenya 110–124 climate change 120– 124 geology 110, 111, 112 stratigraphy 113 sandstone Samburu Hills 110–112, 117, 118, 119– 120 Siwalik Group 160, 161, 172 Tarim Basin 48, 69, 72 sea ice North Pacific 265, 266, 267, 272, 274, 276 influence on winter monsoon 269–271 seaways see Central American Seaway; Indonesian Seaway; Panama Isthmus sediment aeolian central Asia 29– 41 Quaternary, Thar Desert 167, 176 Tarim Basin 46–65, 67–77 fluvial Quaternary Ganga plains 162–165, 174– 175 western India 165– 167 Samburu Hills 114, 117, 118 Siwalik Group 154–177 South China Sea basins 231– 240 Tarim Basin 46, 48, 49–50, 72, 74, 75–76 lacustrine Quaternary, Thar Desert and Ganga plains 168–171 Samburu Hills 114, 115, 117, 119, 121 volcanic, Samburu Hills 110, 112, 113, 114 seismic profiles, South China Sea 222, 223, 224– 230 Shiguanzhai section, OSL dating 91–94, 96, 98, 102, 103 Siberian High pressure system 103– 104 Sikouzi fluviolacustrine sediments 32, 256 silt, aeolian, Central Asia 29– 41

INDEX siltstone, aeolian Tarim Basin 46, 48, 69, 72, 77 grain size analysis 57– 59 linear sedimentation rates 54, 56, 57, 60– 62 provenance 59 Siwalik Group geology 155– 156 Indian summer monsoon fluctuation 154– 155 multi-proxy reconstruction 159– 161, 171– 174 oxygen isotope data 159–161, 171–172, 173– 174 palaeosols 158 –159, 173 stratigraphy 156–158 Song Hong-Yinggehai Basin 220– 221 sedimentation 231, 232, 233, 234, 235, 238, 239 seismic profiles 222, 223, 224–227, 229, 230 structure and tectonics 223, 224– 227, 229, 230, 238 South China Sea East Asian Summer Monsoon evolution, geochemical records 245–260 East Asian Winter Monsoon 10– 11 monsoon reconstructions 221–222, 233, 237, 240 opening 221 productivity 9, 10–11, 13 sediment cores ODP Site 1143 7, 8 ODP Site 1146 221, 234, 236, 237, 246– 260 ODP Site 1148 7, 221– 222, 236, 237, 252, 254 sedimentary records 219– 240 SPECMAP curve 88, 102 speleothem records 90, 100–102, 106 Sphaeroidina bulloides 20, 21, 22, 24 strontium isotope analysis Himalayan– Tibetan Plateau rivers 135, 138, 141, 142, 143 Indus basin sediments 188, 191, 192, 199, 201 Suguta Valley 110 Surai Khola section 157, 158, 159 oxygen isotope data 160, 161, 174 Tadjik Sea 72 Taklimakan Desert 47, 68 aeolian sediment 46, 47–48, 58, 68 formation and evolution 76– 77 Tarim Basin alluvial fans 47, 61–62 Cenozoic sediments 68– 77, 73 desertification 46–63 dust storms 47–48 geology 46–48, 47, 68– 69 tectonics, and the monsoon Greater Himalaya 211, 213 South China Sea 238 –240 temperature sea surface Indian Ocean 18 mountain uplift model 297– 298 North Pacific 271, 274, 276 South China Sea 10–11, 12 Tibetan Plateau uplift modelling 283, 284, 287–288 western Pacific 5, 6 surface air North Pacific 270, 272 Tibetan Plateau uplift modelling 284

307

Thar Desert calcrete 167 Holocene monsoon intensity 177 Quaternary aeolian records 167, 176 Quaternary lacustrine records 168–171 Thatta sediments 187, 188, 189, 190 40 Ar/39Ar mica dating 191, 197– 198, 206, 207 U– Pb zircon dating 208, 209 Tian Shan Mountains 47 loess deposits 46, 47 Tibetan Plateau CO2 uptake 129– 149 chemical weathering rates 140–141, 143 uplift 46, 67–68, 129 Asian aridification 29– 41, 45, 257 East Asian winter monsoon 83, 258, 259 gorge incision 239– 240 impact on palaeoclimate proxies, modelling 279– 289 and monsoon intensification 279 see also Himalaya–Tibetan Plateau tilting, tectonic, Tarim Basin 52, 54, 55, 61–62 Tirr Tirr Formation 110, 112, 113 Tugen Hills 111, 121– 122, 123 U –Pb zircon dating, Indus basin sediments 191, 198, 205–207, 208, 209 upwelling, and monsoon intensity 153 Arabian Sea 153–154 Uvigerina proboscidea 20, 21, 22, 24–25 vegetation, Tibetan Plateau uplift modelling 283–284, 285, 286 Vietnam, uplift 240 Walker Circulation 8, 18 weathering Pearl River Basin 233–234, 236, 237, 238, 239 silicate and CO2 uptake, Himalayan –Tibetan Plateau 129– 149 ODP site 1146 246– 260 winds Central American Seaway 270– 271, 273, 275, 276 Central Asia 31–32 Maldives 18, 19, 25 and monsoon intensity 153 mountain uplift model 294–295 Tarim Basin 47– 48 Wuqia Group 46, 48, 49, 69 lithostratigraphy 70, 71, 72, 73, 75 Wusu Formation 54, 69 Xifeng section, OSL dating 91, 92, 93, 94, 96, 98, 99 Xining aeolian deposit 32, 33, 36, 39 Xiyu Formation 46, 49, 50– 52, 54, 69 age model 52, 53, 54 grain size analysis 57–59 linear sedimentation rates 54, 56, 57, 61 lithostratigraphy 70, 71, 72, 75–77 provenance 59, 60–61 tilting 54, 55 Xunyi section, OSL dating 91, 92, 93, 94, 96, 100, 101

308

INDEX

Yamuna River 131, 133 CO2 consumption 144, 147 silicate weathering 132, 133, 136, 139, 141, 144, 146 strontium isotopes 141, 142, 143, 145 Yangste River see Chang Jiang River Yecheng section 68 aeolian deposits grain size analysis 57–59 provenance 59, 60–61 Cenozoic sediments 69, 73 facies associations 69, 71–72

linear sedimentation rates 54, 56– 57, 60–62 lithostratigraphy 48–52, 69, 70, 72, 73, 75 magnetostratigraphy 69, 70 tilting history 52, 54, 55, 61– 62 Yellow River see Huang He River Yunnan, uplift 239–240 zircons dating, Indus basin sediments 191, 196, 198, 203– 210 Greater Himalaya 205– 207, 208

The Earth’s climate varies through geological time as a result of external, orbital processes, as well as the positions of continents, growth of mountains and the opening and closure of oceanic gateways. Climate modelling suggests that the intensity of the Asian monsoon should correlate, at least in part, with the uplift history of the Tibetan Plateau and the Himalaya, as well as the evolution of gateways and the retreat of shallow seas in Central Asia. Long-term reconstructions of both mountain building and monsoon activity are key to testing the proposed links. This collection of papers presents a series of new studies documenting the variations of the Asian monsoon on orbital and tectonic timescales, together with the impact this has had on environmental conditions. The issue of which proxies are best suited to measuring monsoons is addressed, as is the effect that the monsoon has had on erosion and the formation of the stratigraphic record both on and offshore.