Quaternary paleoenvironmental change on the Tibetan Plateau and adjacent areas (Western China and Western Mongolia)

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Quaternary International 65/66 (2000) 121}145

Quaternary paleoenvironmental change on the Tibetan Plateau and adjacent areas (Western China and Western Mongolia) Frank Lehmkuhl*, Frank Haselein Geographisches Institut der Universita( t Go( ttingen, Goldschmidtstr. 5, D-37077 Go( ttingen, Germany

Abstract This paper summarises the nature of climatic change during the Last Glacial}Interglacial cycle on the Tibetan Plateau and adjacent areas. The results are derived from inland glacier #uctuations, lake level changes and dust (loess) records of Central Asia including Tibet. These are based on our "eld investigations from 1988 to 1998 and literature surveys. The changing extents of di!erent climatic controlled geomorphic landscapes and vegetation zones help to provide estimates of the magnitude of climatic change. Features, such as ice wedge casts, loess, and palaeosols, ELA-reconstruction's and palaeobotanical data are used to help reconstruct palaeoprecipitation and palaeotemperatures. Weathering characteristics, the overlying strata and some luminescence dates indicate that there are two main glacial ice advances during the last glacial cycle. These correspond to marine Oxygen Isotope Stages (OIS) 2 and 4. In some areas, as the eastern or northern margin of the Tibetan Plateau, remnants of older glaciations are preserved. Higher lake levels in the deserts of Central Asia and on the Tibetan Plateau are dated to '40 to 25 ka (OIS 3) and to Late Glacial / Early to Mid Holocene periods. Our work supports the view that in many areas of Central Asia cold phases during the last glacial correspond with the maximum extent of glaciers and the periglacial activity. The last glaciation produced large alluvial fans alternating with periods of high lake stands on the Tibetan Plateau. However, at the northern margin of the Plateau in the Qaidam Basin and in some particular desert areas in Western China, high lake levels occurred also during the Pleistocene and are related with alluvial fans.  2000 Elsevier Science Ltd and INQUA. All rights reserved.

1. Introduction Studies in the central Tibetan Plateau have focused mainly on the distribution and timing of Pleistocene glaciations, lake level changes, palaeoclimatic reconstruction using loess and loess-like sediments. This paper will examine the nature of these studies concerning on the distribution of Pleistocene terminal moraines and the extent of ice on the Tibetan Plateau. From a climatologic and meteorologic point of view the extent of past lakes and glaciers are important, because they provide palaeoclimatic data. The extent of glaciation on the Tibetan Plateau would have in#uenced the distribution of high pressure and low pressure cells, and therefore controlled their tracks eastwards from the Atlantic and regional precipitation. In addition, large lakes in the Qaidam Basin and the Gobi desert must have a!ected the local humidity and autochthonous precipitation.

* Corresponding author. Fax: ##49-551-398006. E-mail address: #[email protected] (F. Lehmkuhl).

This paper also summarises the major research on the drainage systems o! the Tibetan Plateau and Central Asia to provide a framework for the palaeoclimatic studies in this area that covers one of the main aims of Working Group 7 of IGCP 415. This paper considers Central Asia, which is de"ned as the interior endoreic (area without outlet, e.g. Machatschek, 1955) drainage region of Asia, and includes, e.g. the Himalayas and Karakoram, the Pamir, the Tien Shan, and the Tibetan Plateau (High Asia). These regions are very diverse and range from extreme continental climatic conditions with steppe, semi-desert and desert regions and well vegetated mountain areas with relatively higher precipitation values (up to more than 1000 mm/a; Lehmkuhl, 1997c). The distribution of the sand and gobi deserts of Central Asia along with the Loess areas is shown in Fig. 1. The spatial distribution of the lake basins and other locations discussed below is given in Fig. 2 and Table 1. The Tibetan plateau itself is located between Karakoram and Himalayan ranges in the south and Altyn Tagh and Kunlun range in the north. It has a mean elevation between 4500 and 5000 m and many isolated

1040-6182/00/$20.00  2000 Elsevier Science Ltd and INQUA. All rights reserved. PII: S 1 0 4 0 - 6 1 8 2 ( 9 9 ) 0 0 0 4 0 - 3

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Fig. 1. Distribution of sand, loess, sand deserts, gravel deserts (Gobi) and palaeolakes in China and Mongolia. (Modi"ed from Lehmkuhl (1997b)).

mountain massifs of more than 6000 to 7000 m a.s.l. and is the highest plateau of signi"cant extent on Earth. It is generally considered that the uplift of the Tibetan Plateau increased the aridity of Central Asia and induced the monsoon system and the development of loess sequences of the Chinese Loess Plateau (e.g. Prell and Kutzbach, 1992; Ruddimann and Kutzbach, 1989; Molnar and England, 1990; Molnar et al., 1993). Meyerho! et al. (1991) compiled all available data (palaeontologic, palaeoecologic, palaeoclimatologic, depositional rates, volcanics, geothermal, radiometric dates and magmatism, structure and geophysical) and concluded that the uplift of the Tibetan Plateau started at the early Miocene and accelerated by the Pleistocene. The semiarid and arid basins (e.g. Tarim Basin, Junggar Basin and Gobi desert) of Central Asia are situated north of the Tibetan Plateau. They are divided by several mountain ranges (e.g. Tien Shan, Altai: see Fig. 1). As mentioned above, the high mountain ranges and especially the Tibetan Plateau play a major role on the climatic system of Asia and the monsoon systems. This in turn even a!ects global climate and global climatic change (BoK hner, 1996; Flohn, 1981, 1987; Murakami, 1987; Molnar and England, 1990). The Asian monsoon

system consists of South- and Southeast-Asian summer monsoon and winter monsoon (Fig. 3, for details see DomroK s and Peng, 1988; BoK hner, 1994; Benn and Owen, 1998). The "rst is driven by warm and moist air mass over the subtropical and tropical Indian Ocean (Indian monsoon) and Paci"c Ocean (Paci"c monsoon). This brings rain and heat towards East Asia and India in summer. The winter monsoon is driven by a cold mass of the Siberian High Pressure System in winter, generating dust storms (mainly in spring time) and long distance dust transport (for a recent review on dust transport, see Derbyshire et al. (1998)). In addition, the mid-latitude westerlies e!ect the northern and western part of Central Asia and the Tibetan Plateau. This brings moisture to the western mountain ranges of Central Asia, e.g. the Karakoram, the Pamir, Tien Shan, and Altai, and in addition into the interior of the continent. The Tibetan Plateau has an important in#uence on the regional and atmospheric circulation and splits the upper westerly winds in winter into a northern and southern branch. The heating of the plateau in summer raises the air temperatures above the zonal mean for the free atmosphere at the same elevation. This increases the pressure gradient that drives the South Asian summer monsoon.

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Fig. 2. Spatial distribution of dated lake level changes and other locations in Central Asia mentioned in the text. Further details are given in Table 1.

The development of a warm high pressure system in the free atmosphere above the southeast sector of the Tibetan Plateau replaces the westerlies by an easterly surface jet stream, which induces a complete change in surface circulation pattern in summer (BoK hner, 1994). Little is published in western literature on the drainage systems of the Tibetan Plateau and early and middle Pleistocene lacustrine deposits. Li et al. (1995) published a volume, which focuses mainly on the uplift of the Plateau, the long-term climate change (since the Miocene) and detailed geomorphological evidence from the catchment area of the Huang He (Yellow River). He also discusses the evidence for Quaternary glaciations, the loess records over the last 150,000 yr, and the impact of plateau uplift. Summarised studies on Late Quaternary lake level #uctuations are given, e.g. by Fang (1991) Qin and Yu (1998), Tarasov and Harrison (1998). Further references

based on detailed studies and a critical discussion of the di!erent opinions is given below. Many publications have focussed on the loess sequences of the Chinese Loess Plateau and their palaeoclimatic implications (e.g. An et al., 1991b; Hovan et al., 1989; Liu et al., 1985). For a recent review, see Derbyshire et al. (1997). In addition, the terrace sequences of Huang He and other river systems can be dated by the overlying loess}palaeosol sequences (Derbyshire et al., 1995; Ding et al., 1995; Porter et al., 1992; Rost, 1997, 1998).

2. Geomorphic evidence for climatic change 2.1. Spatial distribution and timing of mountain glaciers The timing and extent of Quaternary glaciations on the Tibetan Plateau and adjacent areas is still in debate

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for most areas. Since the beginning of this century there have been heated debates on the extent of a Tibetan ice sheet. (Kuhle, 1988, 1998) has proposed a theory implying the existence of a widespread ice cover of 2.4 million km (2 to 2.5 km thick) over nearly the whole Tibetan Plateau during the di!erent Pleistocene glaciations. In addition, Han (1991) suggested an ice sheet on the Tibetan Plateau in the early Quaternary. Many scientists disagree with this idea (e.g. Derbyshire et al., 1991; HoK vermann et al., 1993a; Li et al., 1991; Lehmkuhl, 1995, 1998a,b; Lehmkuhl et al., 1998; Pu, 1991; Shi et al., 1992;

Table 1(a) Latitude and longitude of dated lake level variations in Central Asia (see Fig. 2) No.

Lake

m a.s.l.

Longitude

Latitude

1 2 3 3 4 5 6 7 8 9 10 11 11 11 12 13 14 15 15 16 17 18 19 19 20 21 22 23 24 25 26 27 28 29 30 31 32 33 34 35 36 37 38 39

N-Tianshuihai Hu Aksayqin Hu Sumxi Co Longmu Co Bangda Lake Tsokar Lake Bangong Co Zhacancaca Lake Zabuye Lake Dagze Co Siling Co Peng Co Pamu Co Nam Co Yamzho Yum Co Da Qaidam Xiao Qaidam Qarhan Dabasun Hu Hsiao Qaidam Co Nag/Xin Xin Qinghai Hu Malingango Zogqen Gaxun Nur/Sogu Nur Badanjilin Shamo Jingshoutu Wuwei Baijian Hu Hoton Nuur Achit Nuur Uvs Nuur/Bajan Nuur Hara Us Nuur Hara Nuur Daba Nuur Dood Nuur Hubsugul Nuur Terkhiin Tsagan Nuur Adagin Cagaan Nuur Orog Nuur Gun Nuur Lop Lake Lake Manas Buir Nuur (outside of map)

4805 4850 5058 5008 4902 4572 4241 4311 4421 4550 4615 4523 4555 4718 4200 c. 2800 c. 3200 c. 2700 c. 2700 c. 3200 4400 3194

79305 79350 80315 80327 81335 78300 79300 82320 84305 87320 88347 90358 90335 90340 90345 95315 95330 95320 95310 96321 98300 100310 99300 98351 101300 101300 101335 102345 104300 88318 90336 92330 92300 93307 98347 99323 100310 99342 100300 100346 106336 90325 86300 117304

35330 35313 34336 34338 34327 33321 33342 32330 31320 31350 31346 31331 31315 30343 28355 37352 37327 36345 37300 36330 35300 36355 32300 32313 42318 42300 42300 38300 39301 48340 49330 50312 47355 47358 48301 51320 50332 48309 45331 45311 50315 40312 45345 47345

c. 3900 850 c. 900 1300 2083 1435 932 1156 1132 2465 1538 1645 2060 1299 1280 600 780 251 585

Table 1(b) Latitude and longitude of selected locations in Central Asia mentioned in the text (see Fig. 2) No. Location C D E F G H I J K L Ice caps A B

Area

m a.s.l.

Longitude

Latitude

Bureghanga Badain Jaran Tengger Shamo Gonghe Basin Dalijia Shan Zoige Basin Nianbaoyeze Shan Chola Shan Yulongxue Shan Kunlun Shan

? ? ? 4645 3400 5369 6120 5596 6626

1043 101320 103330 100347 102341 1023 101310 99300 1003 82302

483 41300 38330 36321 35331 343 33310 32300 273 36351

Gulia Ice Cap Dunde Ice Cap

6710 5325

81329 96324

35317 38306

Zheng, 1989; Zheng and Rutter, 1998; Zhou and Li, 1998). Upon melting such an ice sheet would have resulted in high isostatic rebound. Kuhle argues that ice was cold based and therefore left only minor traces in the "eld and did not deposit any signi"cant amounts of till. In addition, Kuhle (1990) uses `ice marginal rampsa to reconstruct ice margins. However, many workers argued that these landforms are alluvial fan deposits and cannot be used to reconstruct former ice margins (Derbyshire et al., 1991; Lehmkuhl, 1995; Lehmkuhl et al., 1998). Most researches, however, agree that glaciations on the Tibetan Plateau were characterised by valley glaciers and not an extensive ice sheet. In addition, some authors suggest a small ice sheet (morainic platform, cf. Li et al., 1991; Shi et al., 1992; Zhou and Li, 1998) of the penultimate glaciation existed in northeast Tibet at the source area of the Huang He. The existence of an older large plateau glaciation in this area can only be proven by further detailed studies (Lehmkuhl and Liu, 1994). The main reason for the di!erent reconstruction of the extent and distribution of Pleistocene glaciations on and around the Tibetan Plateau is the present of a uniform stratigraphy and nomenclature for the di!erent end moraine sequences or till deposits in Tibet and Central Asia (Lehmkuhl, 1998a,b). Some relative glacier chronologies exist in several mountain regions, but the timing is poorly understood due to the lack of absolute dating. Even a uniform relative stratigraphy based on ELA-depressions as, for example, in the European Alps (Penck and BruK ckner, 1909; for a recent review, see Maisch, 1982, 1995) is not available. Benn and Owen (1998) summarised most of the dating that constrains the timing of glaciations in the Himalayas and in Tibet. Chinese scientists, who argue for a limited extent of the last glaciation on the Tibetan Plateau, show that the Last Glacial Maximum (LGM) ended c. 15 ka

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Fig. 3. Main wind directions in Central Asia and adjacent areas: (a) summer situation, (b) winter situation.

(e.g. Li and Shi, 1992) or c. 13 ka (Li and Pan, 1989). In the Chinese literature the last glaciation is often divided into two stages (approximately equivalent to marine oxygen isotope stages OIS 2 and 4). These were interrupted by an interstadial that lasted from about 55 to 32 ka (e.g. Li and Pan, 1989; Liu et al., 1985; Zhang et al., 1991). Maximum glacier advances are dated for both stages in the various mountain areas (for further details see Benn and Owen, 1998; Lehmkuhl, 1995; Owen et al., 1997a). In addition, for some areas at the eastern and northern margin of the Tibetan Plateau, remnants of Mid-Pleistocene glaciations (penultimate glaciation and older?) are preserved. HoK vermann and HoK vermann (1991) reported on two Pleistocene Glaciations at the northern Kunlun Shan (Fig. 2: L). Li and Pan (1989) and Lehmkuhl (1995) divided three Pleistocene glaciations in the Dalijia Shan, and Nianbaoyeze Shan, respectively (Fig. 2: G, I). Two Pleistocene glaciations could be found at the Chola Shan (Lehmkuhl, 1995; Fig. 2: J) and at the Yulongxue Shan (Ives and Zhang, 1991; Fig. 2: K). Only a few results have been published on Late Glacial and Holocene #uctuations of mountain glaciers on the Tibetan Plateau and the bordering mountains (Lehmkuhl, 1997a; Lehmkuhl and Lang, 2000). Late Glacial Glacier advances are dated to around 15 ka in West-Kunlun, in the Tien Shan and in the mountain areas surrounding the Qaidam Basin (Zheng et al., 1990; Li and Shi, 1992; Yang, 1991; HoK vermann, 1998). Di!erent Chinese authors (e.g. Wang and Fan, 1987; Pu, 1991 argue for several Holocene glacier advances, but most of them separate no further Late Glacial ice marginal margins. In addition, most radiocarbon dating in Tibet has been undertaken on organic deposits that overlay terminal moraines. These give minimum dates and consequently, only a few Pleistocene and Holocene glacier advances are dated with any degree of con"dence (for further discussions see Lehmkuhl, 1997a). During the early Holocene the glaciers were approximately the same size as today. The early Holocene peat

bog in the Kakitu (Qilian Shan, north of the Qaidam Basin and neighbouring to the Dunde ice cap) con"rms this as lying close to the modern glaciers. The peat began to develop at about 9.4 ka (Beug, 1987) and other peat bogs on the Tibetan Plateau started developing about 10 ka (e.g. Lehmkuhl, 1995; Wang and Fan, 1987). HoK vermann (1998) suggest that the coarse sediments in the Qaidam Basin represent small glacier advances at about 9 and 4.5 ka. Historical records, e.g. since the so-called `documental perioda (1100 BC to 1400 AD) for northern China, suggest brief cold and wet intervals during the Late Holocene (Chu, 1973; Zhang, 1991). For example, there are reports on wetter periods in the Taklimagan desert at about 2 ka and during the Little Ice Age (LIA, cf. Yang, 1991). In addition, Wang and Fan (1987) provided evidence for glacier advances in the Neoglacial (or MidNeoglacial) since 4}3 ka and Little Ice Age (16th to 19th centuries). Evidence for these glacier advances comes from wood and branches that are dated in moraines in southern Tibet (Arza glacier, cf. Wang and Fan, 1987) and investigations from other regions, e.g. the Qilian Shan (for details see Pu, 1991). Pollen records showed a cooler (and moisture) period during the Late Holocene (e.g. SchluK tz, 1995, 1999; Sun and Chen, 1991). Furthermore, a younger Holocene temperature depression of about 1}23C at around 3}1.5 ka and for the Little Ice Age can be derived from buried soils in the alpine meadows in eastern Tibet (Lehmkuhl, 1995) and in the Khangai (Lehmkuhl and Lang, 2000) indicating increasing soli#uction activity. 2.2. Pleistocene periglacial features and their implications for palaeoclimate reconstruction Periglacial features, e.g. ice wedge casts, cryoturbations, and soli#uction layers, can provide evidence for palaeotemperatures (see Karte, 1983, 1990; Lowe and Walker, 1997). Further palaeoclimatic evidence can be

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Table 2 Climatic thresholds and climatic-diagnostic values for di!erent periglacial features (modi"ed according to Karte, 1983, 1990). MAAT"mean annual air temperature, MAT"mean air temperature (month) Periglacial features

MAAT (3C)

MAT (3C) coldest month

Annual precipitation (mm)

Other climatic implications

Ice wedge polygons

(!4}!8

(!20

'50 } 500

Sand wedge polygons

(!12}(!20

Micro frost crack polygons (Molisolfrostkeile)

(0}(!4

Rapid decrease in temperature in early winter time, thin snow cover Rapid decrease in temperature in cold and arid environments Rapid decrease in temperature in temperate * continental environments

(100 (!8

derived from the palaeosols sediment cover and aeolian mantles on bedrock or terraces. For example, in eastern Tibet (Basin of Zoige, 3400 to 3500 m a.s.l.; see Fig. 2) a horizon with fossil involutions (cryoturbation) and an ice-wedge cast "lled with sandy loess provide evidence of former permafrost conditions (Lehmkuhl, 1995; Lehmkuhl and HoK vermann, 1996). A layer of homogeneous sandy silt approximately 140 cm thick covers this horizon. The basal layers of this silt have been dated by thermoluminescence (TL) to be c. 18 ka (Lehmkuhl, 1995). It is generally accepted, that permafrost will only occur where the mean annual air temperature (MAAT) drops below 03C, probably 4!23C in mountain regions (cf. Lowe and Walker, 1997; Washburn, 1979). However, according to di!erent authors (summarised in Karte, 1983, 1990; Table 2), !63C is accepted as the maximum MAAT possible for development of ice-wedge casts. Based on the present MAAT within the Zoige Basin (MAAT of Zoige: 0.73C) and the lowering of permafrost from the current lowermost boundary in 4300 m a.s.l. to at least 3300 m a.s.l. during the Last Glacial, the inferred temperature decrease was 563C (Lehmkuhl, 1995). This reconstruction of palaeotemperatures and the existence of ice wedges suggest higher aridity (see Table 2). This "ts well with the snowline depression of the last glaciation within this area that Lehmkuhl and Liu (1994) and Lehmkuhl (1995) calculated to be 600}800 m. In the source area of the Huang He, close to the settlement of Madoi and the Xin Xin Lake (Fig. 2, No. 17), two layers of soli#uction debris at the slopes can also be found. At the top of the hill a section contains sand wedges that are 42 m deep. These have TL dates of 18 ka equivalent to the global LGM (further details are published in Lehmkuhl, 1995; Lehmkuhl et al., 1998). However, involutions and single sand wedges may form outside of permafrost regions. For example, Sun and Li (1986) describe such features in the Datong Basin, in the lowlands of China (c. 1000 m a.s.l.; 40306N/113320E). These features did not necessarily require permafrost conditions (Rost, 1998), but according

?

to the photos given by Sun and Li (1986) they could be remnants of tree roots. However, they were used in estimating the palaeotemperatures of permafrost environments in Central China. This may lead to the relatively high deviations of annual mean temperature for the maximum cooling of the last glaciation with values of '103C colder than the present-day mean temperatures (Frenzel et al., 1992). 2.3. Fluctuations of glaciations and lake level: an example form the Nam Co, Southern-Central Tibet The southern margin of the Nam Co (or Tengri Nur; Fig. 2, No. 11) will be described in detail to provide an example for the di!erent timing of lake level high stands and glacier advances on the Tibetan Plateau. Field work in this area provides evidence for several former lake levels #uctuate and two major glaciations (Fig. 4). The maximum extent of the oldest glaciation in this region occurred near the highest peaks of the Nyainqe( ntanglha Shan (7162 m). These glaciers reached down to the present day lake level (Fig. 4, No. 1). However, there are no lakeshores that relate to that ice margin. After this glacier advance, a high lake shoreline partly destroyed these oldest moraines forming a cli! about 30 m above the modern lake level (Fig. 4, No. 2). Next a younger glacial stage terminated upstream of the oldest terminal moraines (Fig. 4, No. 3). The #uvial outwash of this stage destroyed parts of the oldest moraines and the cli!. Seven former lake shorelines (VII to I), up to 20 m above the modern lake level, have been mapped on top of the youngest glacio#uvial outwash (Fig. 4, No. 4) providing evidence for the youngest recession of the lake. Holocene aeolian sandy silt has been accumulated on top of the highest shorelines and tills and up to the upper limit of vegetation that is presently at about 5500 m a.s.l..

 Co"lake in Tibetan language, Nur or Nuur"lake in Mongolian language.

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127

Fig. 4. Landscape development at the lake Namco and the northern slope of the Nyainqe( ntanglha Shan. Proposed time periods are given in Table 3. The roman numbers indicate the di!erent lake shore lines. Further explanations see text.

Table 3 Proposed time frame for di!erent phases of the relief development at Lake Nam Co and northern slope of the Nyainqe( ntanglha Shan (Central Tibet) Morphology and sedimentology

Period

1. First glacier advance: Accumulation of older morainic and alluvial fan deposits 2. Transgression of the lake: Erosion of a cli! or highest beach deposits (VII) 3. Second main glacier advance: Accumulation of younger morainic and alluvial fan deposits 4. Transgression (beach wall VI) 5. Regression and reduction of #uvial activity (beach walls V}I)

Early stage of Last Glaciation OIS 4 (70}50/40 ka) Interstadial of Last Glaciation OIS 3 (40}30 ka) Latest stage of Last Glaciation OIS 2 (32}18/13 ka) Late glacial (since 13 ka) Late glacial/Holocene OIS 2 and 1 (13}8 / 5ka)

See Fig. 3 and text for further details.

Weathering characteristics and the overlying strata indicate that the youngest glacial stage should be equivalent to the last glacial cycle (corresponding to the OIS 2; Table 3). The traces of higher lake levels should be equivalent to the Late Glacial/Early to Mid Holocene in

age when wetter climatic conditions persisted all over Tibet and Central Asia (see below). The older moraine probably corresponds to the early stage of the last glaciation (equivalent to OIS 4) or to the penultimate glaciation (OIS 6). The highest, 30 m high, shoreline may be

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of last Interstadial age (equivalent to OIS 3) when numerous lakes in Tibet and Central Asia have high stands (see below). This example shows clearly, that lacustrine phases (or periods with high lake levels) alternate with glacier advances at least in this area of the Tibetan Plateau.

3. Palaeoclimatic implications of Late Quaternary dust sources in Central Asia Dust deposits, mainly clay- and silt-sized particles (H (0.6 mm), which formed during the Quaternary, are the result of a combination of processes (e.g. Pye, 1987). These include four main sets of proxies. The "rst involves the production of dust by weathering of host sediments with a high content of clay and silt. Pye (1987) recognises at least 11 di!erent types of terrain that are favourable for dust de#ation. Of interest here are the glacial outwash plains, alluvial fans, wadi sediments, lake and playa sediments and stony desert (gravel gobi). The second involves the source area, where the surface conditions are favourable for the de#ation of dust. The subtropical desert regions of North Africa and the dry high altitude basins of Central Asia are suggested as the major sources of present-day dust de#ation (Pye, 1987; Duce et al., 1991; Zhang et al., 1997). The third involves the transport of dust where atmospheric conditions are favourable for de#ation and transport. The deserts of Central Asia and China are considered to be the major dust source region of present day long distance dust transport. Dust is transported as far as, e.g. Japan, Hawaii or Alaska (Goudie, 1983; Pye, 1987; for a recent review, see Derbyshire et al., 1998). Quaternary dust from Central Asia also a!ects the surrounding Oceans and can be found in the deep-sea cores from the Arabian Sea (e.g. Clemens and Prell, 1990, 1991; Sirocko et al., 1993, 1996)

and the Paci"c Ocean (e.g. Rea, 1994; Rea and Leinen, 1988; Hovan et al., 1989, 1991). The global input reaches up as far as Greenland and was found in the GSIP2 ice core (Ram and Koenig, 1997; Biscaye et al., 1997). The "nal set involves the deposition of dust where atmospheric and surface conditions are favourable in order to allow long-term sedimentation of dust (Table 4). In Central Asia and adjacent areas dust deposits mainly occur as loess and loess-like sediments in various environments of China and Mongolia including the Loess Plateau of China (Fig. 1). However, in the high mountains areas of Tibetan Plateau thin layers of loesslike sediments also are observed up to 5300}5500 m a.s.l. (Lehmkuhl, 1997b). In the lacustrine environments of the Qaidam Basin (913}983E; 363}38330N) loessic silts are associated with halite deposits (Chen and Bowler, 1986) and silt-sized particles have also been found in the lacustrine sediments of Qinghai Hu (100310E; 36355N; Kelts et al., 1989; Lister et al., 1991). In addition, glacial environments also show accumulation of dust, e.g. in the ice core drillings from Dunde ice cap (Fig. 2: 96324E; 38306N; Thompson et al., 1989) and Guliya ice cap (Fig. 2: 81329E; 35317N; Thompson et al., 1997). A striking feature of most dust records (terrestrial and marine) is the enhanced dust #uxes indicated by higher sedimentation rates (by an average factor of 20) during glacial times compared with the interglacial values (Mahowald et al., 1999). For example, the Vostok dust record shows the lowest dust concentration during the interglacial times. With respect to the Holocene values, the dust #ux was 15 times higher during OIS 2 and 11 times higher during OIS 4, respectively (Petit et al., 1990). Modelling of LGM dust dynamics suggests a three times higher total dust emission in LGM with respect to modern values (Mahowald et al., 1999). Measurements of magnetic susceptibility of loess sequences from the Chinese Loess Plateau (e.g. An et al., 1991a, c; Hovan

Table 4 Proxy indicators, units, and chronological methods of dust records according to di!erent authors Authors

Proxy

Unit

Chronology

De Angelis et al. (1987) Jouzel et al. (1989) (Vostok ice-core) Jouzel et al. (1993) (Vostok ice-core) Kukla et al. (1988) An et al. (1991a) (Loess Plateau, China) Petit et al. (1990) (Vostok ice-core) Rea and Leinen (1988) (Paci"c cores) Sirocko et al. (1996) (Arabian sea-core)

Al-pro"le (as indicator of terrestrial input) Dust content variations Loess

ng/g

Glaciologic model

cm/g Magnetic susceptibility

Dust#ux Aeolian component Terrigenous matter

mg/cm/ka mg/cm/ka %

Thompson et al. (1989) (Dunde ice-core, China) Thompson et al. (1997) (Guliya ice-core, China)

Dust concentration

1 ka averages of total particles

Dust concentration

Number of particles/ml

Glaciologic model C, TL, magnetic reversal SPECMAP Glaciologic Oxygen isotopes AMS-C chronology  (interpolated) Ice-core chronology (extrapolated) dO correlation with CH concentration of GISP2 and SPECMAP

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et al. 1989) show lowest values during Glacial times (OIS 2 and OIS 4) and this indicates weak pedogenesis and enhanced aeolian deposition. Although Phillips et al. (1993) noted that: `2 previously the large aeolian dust #uxes out of Central Asia during Glacial maxima have been considered to be associated with increased aridity in the region2a. It can be shown that, with respect to the above-mentioned processes which lead to dust deposition, the explanation for increasing dust deposition rates or continental aerosol input vary according to di!erent authors. Those include: increasing aridity in source areas (Sirocko et al., 1996; Rea and Leinen, 1988); expansion of arid areas (Jouzel et al., 1989; Kolla et al., 1979; Petit et al., 1981; De Angelis et al., 1987); changes in vegetation (Mahowald et al., 1999); occurrence of new source areas, including exposed continental shelves owing to sea-level lowering during glacial intervals (Petit et al., 1981; Petit et al., 1990; Ono and Naruse, 1997); and changing atmospheric conditions including monsoon dynamics (Feng et al., 1998); higher wind speeds (Petit et al., 1981; An et al., 1991a); more e!ective atmospheric transport (Jouzel et al., 1989); and persistence of strong winds (Phillips et al., 1993). However, a combination of two or more of these processes is proposed by most authors (e.g. Petit et al., 1981, 1990). In the following it is shown that since 32 ka climatic induced landscape change could be another potential explanation of enhanced dust emission in Central Asia.

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Fig. 5. Late Quaternary dust records and palaeo-precipitation of the Qaidam Basin. (A) Terrigenous content of the core 74KL (Arabian Sea; Sirocko et al., 1996); 1"AMS uncalibrated 14C-dates; 2"calibrated radiocarbon-dates. (B) Dust #ux of the Vostok ice core (Antarctica; Petit et al., 1990). (C) Dust concentrations of the Dunde Icecap (Tibetan Plateau; Thompson et al., 1989). (D) Palaeo-precipitation of the Qaidam Basin (Qinghai, China; HoK vermann, 1998). 1"lake level; 2"precipitation.

4. Climatic induced landscape change since 32 ka in Qaidam Basin Qaidam Basin is located at the northern border of the Tibetan Plateau in Qinghai Province, China (see Figs. 1 and 2). Mean elevation of the basin #oor is &2700 m a.s.l., whereas the surrounding mountains of Kunlun-, Altyn- and Qilian Shan reach elevations above 5000 m a.s.l.. Mean annual precipitation is between 25 and 50 mm and mean annual potential evaporation is &3000 mm. Prevailing winds are westerly and northwesterly. The terrain is characterised by salt lakes, playas, and yardangs. Aeolian processes a!ect the landforms. The ice core from the Dunde ice cap, located in the Qilian Shan at the northern border of the Qaidam Basin (Fig. 2), provides a dust record spanning the past 40 ka (Fig. 5C, Thompson et al., 1989). According to these authors, the dust deposited in the ice cap originates from the surrounding deserts. HoK vermann (1985, 1987) recognized di!erent arid and semi-arid landscape types or landscape regions in Africa and Central Asia. Besides the bedrock geology and tectonics, the climate infers di!erent geomorphic processes and vegetation, which have an e!ect on the landforms. Based on several "eld studies, "ve landscape regions can be

separated in the arid areas of Central Asia (HoK vermann, 1985, 1987; HoK vermann et al., 1993; Lehmkuhl, 1997c). Table 5 shows the di!erent landscape regions including the main processes, vegetation, and climatic implications for the semiarid and arid regions of Central Asia. These di!erent landscape regions changed in this area during the Pleistocene. For example, throughout the cold phases of the Pleistocene large alluvial fans and fanglomerates or pediments (bajadas) were deposited in the basins and at the foothills of the mountains due to larger sediment supply in the catchment areas. These alluvial fans (or pediments) extended several kilometres beyond the mountain front, in some areas more than 40 km. In addition, these landforms produced large quantities of silt and sand during the Pleistocene (Derbyshire et al., 1998; Lehmkuhl, 1997b). Since the beginning of the Holocene, however, there has been a reduction in sediment supply due to increased vegetation cover and less cryogenic weathering. The modern rivers and streams incise into their Pleistocene gravel and fanglomerates. In the basins they create smaller sized alluvial fans. This fan trenching has become dominant due to the reduction of sediment load and concentration of precipitation in more

Widespread alluvial fans, fanglomerates, and pediments (bajadas); braided rivulets Development or preservation of planar surfaces Isolated gullies with periodic discharge Extended gully erosion with stepped valley pro"les; weak slope processes and soil formation Steep #uvial erosion forms, mainly gorges Strong structural control Plains of various amounts of sand or di!erent sand cover on bedrock Development of sand surfaces (and dune"elds) Yardangs and dune"elds Forming of dune"elds, wind-formed features including yardangs and de#ation hollows

Pediment region (Fu{#aK chenlandschaft)

Steppe gorge region (Steppenschluchtenlandschaft)

Desert gorge region (WuK stenschluchtenlandschaft)

Desert plain region (Sandschwemmebenenlandschaft)

Wind formed region (Aerodynamische Landschaft)

Aeolian processes +IV : f3a

Aeolian and episodic ally #uvial (denudation) processes, periodically intensive sheet #ows +IV : f3a (s1)

Intensive episodically erosion processes +IV : f3

Weak slope wash and slope erosion, weak #uvial incision +IV : f2/f3s1a

Slope wash, strong episodic #uvial processes +IV : f3s1a

Main processes regions with di!erent combination of geomorphic processes (Hagedorn and Poser, 1974)

Desert

Desert

Semi-desert with saisonal and patchy grass or steppe vegetation

Steppe

Desert steppe with bushes and episodic grasslands

Vegetation

Hyperarid (30 mm/a

Hyperarid 30 bis 50 mm/a

Arid 50 to 150 mm/a

Temperate}semiarid '150 mm/a, ¹ (!23C (

Continental}semiarid 150}350 mm/a, frost weathering

Climatic implications

Note: These FPR are associated with distinct geomorphological processes (HoK vermann, 1985, 1987; Lehmkuhl, 1997c) and units of vegetation. Di!erent climatic conditions can be derived for these regions.

Main landforms/main forming processes

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Landscape region (Formungs- und Proze{region: FPR)

Table 5 Characteristics of di!erent landscape regions (climatic-induced types of landforms"Formungs- und Proze{regionen"FPR)

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single rainstorm events in vast areas of Central Asia (HoK vermann, 1985, 1987; HoK vermann et al., 1993b; Lehmkuhl, 1997c). Whereas in the semi-arid and arid regions the incision of small gullies and also gorges occur (steppe gorge region or desert gorge region, see Table 5), the hyperarid area is characterised by aeolian processes (desert plain region and wind formed region, see Table 5). Based on the modern climatic implication of the di!erent landscape regions it is also possible to estimate past precipitation and temperature. Furthermore, HoK vermann (1998) calculated the annual precipitation values of Qaidam region since 32 ka by comparing the annual sedimentation rate of clastic sediments in core CK 2022 from the central Qaidam Basin (94352E, 37305N, Huang et al., 1981; HoK vermann and SuK ssenberger, 1986; see Fig. 5D). With regard to the Pleistocene dust #ux, there is obviously no strong correlation between the dust deposition and mean annual precipitation. Nevertheless, Goudie (1983) points out that the peak of dust storm frequency occurs in areas where rainfall ranges between 100 and 200 mm/a (Fig. 6). This is in accordance with our own more regional observations made in Tibet, Mongolia, and Africa (HoK vermann, 1985, 1987; HoK vermann et al., 1993a,b; Lehmkuhl, 1997b,c). However, as already Goudie (1983) notes: `The relationship between dust storm frequencies and mean annual precipitation is not a simple one.a Fig. 7 shows a model of Late Pleistocene dust deposition #uctuations in the Qaidam Basin including precipitation values, dust records, lake level variations and landscape evolution.

Fig. 7. Proposed schematic model of Late Pleistocene dust dynamics in the Qaidam region. Size of arrows is tentative and ages in BP. Further explanations see text.

Fig. 6. Frequency of dust storms in relation to precipitation (Goudie, 1983) and geomorphic landscape types according to HoK vermann (1985).

The interval between 32 and 24 ka and probably some time before 32 ka is characterised by advancing glaciers and rising lake levels in Qaidam Basin. HoK vermann and SuK ssenberger (1986) named the time beginning at the last glacial cycle with wet and cold conditions the `Anaglaciala and the following dry and cold phase (24 to 15 ka) the `Kataglaciala. Chen and Bowler (1986) found sedimentary evidence for the occurrence of a freshwater to slightly saline lake. This formed under more humid conditions in Qarhan area between 40 and 25 ka (the `pre-evaporitic stagea) comparing with the present day

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condition. For the same area, Liu (1991) reported on the occurrence of a shell bed at Qarhan salt lake indicating freshwater conditions and very humid climate from 38.6 to 28.7 ka. Since 25 ka the lakes began to desiccate. Huang (1994) published dates of between 31.8 and 23.8 ka from higher lakeshores (#8 to #13 m) of the Xiao Qaidam. In addition, palaeobotanical evidence shows that humid environmental conditions existed between 31 and 24 ka (Kong, 1993). Precipitation during this humid period was between 150 and 350 mm/a (Fig. 5D). This is the particular precipitation range for the landscape region of the pediments. As argued above, under the present day climatic conditions, the pediments and large alluvial fans are inactive in most parts. Due to #uvial erosion they became dissected mainly in the upper parts, while dunes and desert plains invade the lower parts. In addition, the Holocene fan size is being reduced to about 10}30% of the Pleistocene one. This is con"rmed by ground checking and the interpretation of satellite images. In addition, it can be shown that during the time of highest lake level (#170 m, 2850 m a.s.l.), around 25 ka (Fig. 5D), the pediment regions have been active and were interlaced with lakeshore features. It is a convincing argument that the period of rising lake level in PalaeoQaidam-Lake was also a period of active pediment expansion (HoK vermann et al., 1993a,b). The dust record of Dunde ice cap (Fig. 5C; Thompson et al., 1989) shows that dust #ux increased to a maximum about 30 ka. During the Anaglacial period (24}15 ka), the glacial and periglacial landscapes of the adjacent mountains were expanded due to a decline of temperature of approximately 63C with respect to modern values (HoK vermann, 1998). This resulted in an increase in amount of clastic material reaching the basin. Since the processes of alluvial fans and fanglomerates (pediment regions) are generally characterised by transport and accumulation of weathered material from higher mountains areas to the basins, the pediment regions are prominent dust source areas in the semiarid environment. The pediment landscape in the Qaidam region is thus interpreted by the authors as a major dust source between '32}24 ka. The interval between 24 and 21 ka is characterised by the decline in the area of pediments and lake desiccation due to reduced rainfall (50}300 mm/a). This is the early stage of the so-called `Kataglaciala (24}15 ka cf. HoK vermann and SuK ssenberger, 1986). This period is characterised by higher temperature values and initiation of glacier retreat. Chen and Bowler (1986) call this the `evaporitic stagea of Palaeo-Qaidam-Lake. The pediments were still acting as dust sources while the shrinking lake pro-

 Golmud in the Qaidam Basin (94306N, 36312E) in 2808 m a.s.l. has a MAAT of 4.23C (January: !10.9, July: 17.6) and an annual precipitation of 36 mm (average 1951 to 1980, cf. DomroK s and Peng, 1988).

vides playas and salt #ats favourable for de#ation. De#ation of dust, therefore, remained high. Lake level lowering was still in progress between 21 and 15 ka. The large areas of exposed lacustrine sediments provide the main dust source. The climate became more arid and remained cool. The precipitation was between 50 and 150 mm/a. Under these climatic conditions, fan entrenchment became dominant. This change in the processes results also in a change of the landform regions from the so-called pediment region towards the landform of desert gorges (Table 5). The latter region is characterised by period or episodic discharge caused by heavy rainfall. It results in fan trenching. However, debris #ows with "ne-grained clastics still reached the basin. Therefore some silt-sized material was still available. In addition, the drainage channels dry out and also became an important source of dust. After &15 ka, hyperarid climatic conditions prevailed. The driest interval of the `evaporite stagea occurred between 15 and 9 ka. This is indicated by the uppermost halite zone at Qarhan section (Chen and Bowler, 1986; Liu, 1991). According to HoK vermann and SuK ssenberger (1986) a cold and dry interval occurred between 14 and 10 ka. This is consistent with evidence for deposition of aeolian silt and loess accumulation around 17}15 ka in Qinghai Basin (`Erlangjiang terracea, Liu, 1991). In addition, sedimentation of sandy facies of possible aeolian origin occurred at Qarhan at 16.4 ka after the desiccation of lakes at Xiao Qaidam and Kunteyi soon after 15 ka (Bowler et al., 1986). A single U/Th date from the western Qaidam Basin indicates initial halite deposition at 16.3$2.2 ka (Phillips et al., 1993). The Hyperarid climate remained until at least 9 ka, and subsequently there were some brief episodes of wetter conditions and alternating cycles of rising lake level and desiccation (cf. Chen and Bowler, 1986). According to Bowler et al. (1986) the present day climate in Qaidam basin represents the driest part during the last 40 ka. According to the calculations of HoK vermann (1998), the precipitation values were generally below 50 mm/a since &14 ka, and two more humid intervals occurred around 9 and 5 ka (Fig. 5D). This hyperarid climate encouraged the development of a desert plain region (Table 5). Since desert plain landscapes have virtually no drainage (HoK vermann, 1998), only few clastics reached the lake basin. This means that less material was provided for de#ation with respect to the period before 14 ka. De#ation of dust would therefore decline. Additional support is given by the dust #ux values of Dunde icecap (Thompson et al., 1989). In the Dunde record, the concentrations of anions (Cl\ and SO\) increased grad ually between 30 and 10 ka, simultaneously with high dust concentrations. However, since 11 ka an inverse trend has taken place: the dust #ux values of Dunde icecap dropped sharply soon after 11 ka while the anion concentrations increased (Thompson et al., 1989).

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A comparable trend is recorded from Guliya ice core (81329E; 35317N; Thompson et al., 1997), although chronological control is somewhat confusing. Thompson et al. (1989) suggest that increase of anion concentrations in the Dunde record may illustrate drying of freshwater lakes and decrease of dust concentrations could re#ect changes in the capability of the atmosphere to transport dust. Winkler and Wang (1993) suggest that strengthened southwestern monsoon and deposition of sea salt aerosol after 11 ka is another possible interpretation of Dunde data. However, using Ockham's razor the assumption of drying freshwater lakes seems more probable and is consistent with other palaeoenvironmental data. The shrinking of dust source regions (mainly glacial, periglacial and pediment regions) and the establishment of desert plain landscape region are the probable explanation for the lowering of dust #ux values in the Qaidam Basin as observed in the Dunde ice core.

5. Palaeoclimatic implications of lake level variations According to calculations based on literature, maps and satellite imagery (Atlas of false colour Landsat Images of China, 1983) the Tibetan Plateau is covered by lakes between 743}983E and 283}403N exceeding approximately 30,000 km (more than 27,000 km after Zheng, 1989; cited by Zhao, 1994). The area of recent day playas, which are assumed to show evidence of higher lake levels in the past or even the existence of a palaeolake, has been calculated to about 20,500 km. This is interpreted as the minimum value of lake level #uctuations. If the evidence of higher lakeshores from "eld investigations, literature and interpretation of satellite images together with palaeolake levels reported by Atlas of Tibet Plateau (1990) are taken into account, the overall area of palaeolakes is estimated to app. 84,000 km. This is interpreted as the maximum value of lake level #uctuations. This poses the question of whether high lake levels on Tibetan Plateau occurred simultaneously or at di!erent times with respect to di!erent regional climatic and environmental evolution. Two di!erent approaches can be used to extract palaeoclimatic information from lacustrine records, which are described in the literature. These include: 1. Quantitative methods, which include studies of lakeshore features, remnants of lacustrine sediments or other indications of lacustrine environment at elevations above, or * less common * below the present lake level. This provides information of lake level elevation and thus extension of lake area (see Frenzel, 1994; HoK vermann and Lehmkuhl, 1994a). 2. Qualitative methods, including sedimentological, mineralogical, palynological, palaeontological and stable isotope studies of pro"les from lake drilling cores. In most cases several methods are applied (e.g. HoK vermann and SuK ssenberger, 1986; Fontes et al., 1996;

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Rhodes et al., 1996). Synoptic studies of both, quantitative and qualitative methods provide detailed information of lake evolution (e.g. Gasse et al., 1991; Van Campo and Gasse, 1993; WuK nnemann et al., 1998). Chinese scientists have studied the development of the Huang He on the eastern fringe of the Tibetan Plateau. They have outlined in detail the river history in relation to the uplift of the Plateau. Several lacustrine deposits of Pliocene and Early Pleistocene age provide evidence for extensive lakes in the catchment area of the Huang He (Li, 1991; Li et al., 1995). Two to three main phases of uplift can be recognised in the Pleistocene and can be correlated with the river history (terraces and Loesspalaeosol sequences near Lanzhou, see Fig. 8). According to Li et al. (1995) the modern Huang He catchment was separated into several independent catchment areas between 3.4 and 1.7 Ma ago and some of them were interconnected in the middle reaches between 1.6 and 1.5 Ma (Qinghai-Xizang Movement, &2000 to 1000 ka; Part 1 of Fig. 8). The Sammen and Jishi Gorge were cut at 1100 ka due to a strong uplift of the Tibetan Plateau (Yellow River Movement), and the upper reaches of the Huang He stretched up to the Xunhua Basin (Part 2 of Fig. 8). The development of Longyang Gorge, associated with the Gonghe Movement, in the middle reaches of the Huang He was dated to about 150 ka (Li et al., 1995). According to Li et al. (1995), the Gonghe Basin (Part 3 of Fig. 8) was reached at 100 ka and, Li (1991) suggested that the upper reaches (the present Huang He drainage) of the Huang He (upstream of Rouergai or Zoige Basin, Part 3 of Fig. 8) is younger than 10 ka. Dates at Jungong (Fig. 8-3 : 37,610$2010 BP) and in the Zoige Basin (22,650$300 BP; see Part 3 of Fig. 8) were taken by Li (1991) as indications of a young age with respect to the geological time scale for the development of the upper course of the Huang He. According to investigations made by the present authors, we can con"rm a lacustrine period in the Zoige Basin around 10 ka. This is based on a radiocarbon date of molluscs: 10,865$305 BP (Hv 20316) and an OSL-date of 9.4$1.0 ka on silt material above (HDS 50, Lehmkuhl and Lang, 2000). However, clear "eld evidence shows that the river course was already established at least since 200 ka and the deposition of lacustrine sediments in the Zoige Basin may have been climate controlled. Upstream of the Zoige Basin several terraces of the Huang He and di!erent land surfaces provide evidence for an older river course (Lehmkuhl and SpoK nemann, 1994). A second terrace could be dated to be of penultimate glaciation (cf. Lehmkuhl, 1995). Fang (1991) gives a general overview on the lake evolution of the last 30 ka in China. The maximum extent of Quaternary lake levels is presented in two Chinese maps (Atlas of Tibet Plateau, 1990; Li et al., 1991). The latter also shows the extent of Pleistocene glaciations on the Tibetan Plateau. Most other papers focus more on a

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Fig. 8. Drainage pattern in the catchment of the Huang He since the last 1.7 mill. yr. (Modi"ed from Li et al., 1995).

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Fig. 9. Lake level variations and arid/humid intervals in Central Asia according to di!erent authors. 1, 3, 4, 5, 6, 8"Fang (1991), 2"Li and Shi (1992), 7"Yan and Petit-Maire (1994), 9, 10"WuK nnemann et al. (1998), 11"Hofmann (1993).

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regional scale. We divide the di!erent lake basins with dated samples or cores into "ve regions. An overview including the names of the lakes is given in Fig. 2 and Table 1. Fig. 9 shows lake level variations and arid/humid intervals since 30 ka in di!erent regions of Central Asia according to various interpretations. It is striking that there is obviously no synchronous humid or arid interval and high or low lake level across the whole plateau. A review and discussion of lake level history in Central Asia with respect to primary literature and our own results is discussed below to provide a synoptic interpretation. 5.1. Western Tibet In the Chinese literature di!erent opinions about the chronology of even regional scaled lake level changes have been published. This may be re#ected by completion of Fang (1991, p. 47): `2On the Qinghai-Tibetan Plateau and alpine regions of western and southwestern China2 lake levels reached their maximum as early as about 15,000 yr BP and/or 12,000 yr BP after the LGM desiccation or de#ation. 2Western China, like other areas of the country, experienced prolonged and intense aridity between 16,000 yr BP and 10,000 yr BP2a According to Wang et al. (1990) high lake levels in western Kunlun area (western Tibetan lakes: Aksayqin, Bangda, Guozha, Tianshuihai and Northern Tianshuihai; Fig. 2, Table 1) occurred around 46, 39}24, 22}18.5, 16}13 and 11}9 ka. However, Li and Shi (1992) report that the lakes experienced higher levels during 40}30, 22}15, 9}5 and 4}2.5 ka. The existence of higher lake levels around 18 ka is a subject of special discussion (Fang, 1991). However, discussion of data from Chinese literature is sometimes di$cult, because of lack of information about dated material and detailed environmental and topographic description (Lehmkuhl, 1997a). A careful interpretation of the data (Wang et al., 1990; Li et al., 1989; Fang, 1991; Li and Shi, 1992) suggests that higher lake levels occurred during '40}24 ka and &16}13 ka. Frenzel (1994), however, interpret the data of Li et al. (1989) in a di!erent way. A review of the dates in the Chinese literature shows that the Chinese dates mainly cover the Late Pleistocene. The chronology of lakes Bangong Co, Sumxi Co and Longmu Co, situated only (200 km away from the above-mentioned lakes, is based mainly on Late Glacial and Holocene dates (Gasse et al., 1991; Fontes et al., 1993; Fontes et al., 1996; Gasse et al., 1996; Van Campo et al., 1996; Fan et al., 1996). Di!erent regional settings may a!ect lake level changes. This includes catchment area, elevation of lake basin, glacier meltwater input etc. This discrepancy also can be explained to some degree by di!erent chronological methods. It must be suggested that dating control cited in the above mentioned Chinese literature is based

on conventional C dates, but no additional information concerning C-dating method is given. Dated material is mainly sediment and aquatic plant, but this is often not reported. The chronology of lakes Sumxi Co and Longmu Co (Gasse et al., 1991; Van Campo and Gasse, 1993) is founded on 14 C dates with oldest +1 ages of 12.7 and 12.5 ka. Fontes et al. (1993) used their data and 21 additional dates from the same lake basins to provide a validity check and age corrections of their samples. Fontes et al. (1993) established a chronological framework for Bangong Co based on 11 C ages +1 corrected from the aging e!ect. It is striking to note that the oldest (uncorrected) samples are 18,999$240 BP from Sumxi Co (Fontes et al., 1993) and 16,110$210 BP from Bangong Co (Fontes et al., 1996) which both are regarded as invalid. For marine samples a subtraction of 400 yr is usually applied to compensate for apparent ageing (Stuiver and Braziunas, 1993). The corrected ages for western Tibetan lakes become remarkably younger and thus the chronological framework too. So it is impossible to compare time scales from lakes dated by conventional C method with chronologies based on corrected C dates of lacustrine carbonates. It seems more probable to compare only the conventional ages under the assumption that synchronous events of lake histories can be identi"ed in spite of problems of absolute dating. In addition, it should be pointed out that most of the lacustrine samples from Central Asia, which are cited in the literature, have been dated by conventional C method. Li et al. (1989) noted that their dated material is mainly of organic origin and only a few derived from calcareous content of sediments. The dates reported from Sumxi Co and Longmu Co (Gasse et al., 1991; Van Campo and Gasse, 1993) are from shells, ostracodes, aquatic plants and algal mats. Samples from Bangong Co include shells, total organic matter, ostracodes, aquatic plants and CaCO , and the uncorrected ages are roughly  consistent between &16 ka (12.30 m) and &5 ka (1.70 m) with respect to age versus depth of pro"le ratio, although two dates from total organic matter show slightly younger ages. Based on the uncorrected ages from Bangong Co (Fontes et al., 1993), the lake and climatic evolution can be interpreted as follows. Before &16.2 ka a closed lake under arid conditions prevailed as can be interpreted from high salt concentrations and high O values induced by high evaporation. The interval &16 to &12.7 ka is characterised by the opening of the hydrological system under more humid climatic conditions with positive precipitation-evaporation ratios. According to

 Fontes and Gasse (1989) report a sample from a living (sic!) snail (Melanopsis sp.) from Wadi el Akarit (Tunesia) which was a!ected by the `hard-water-e!ecta, providing a radiocarbon age of 13 ka BP!

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Fontes et al. (1993), this humid interval is due to enhanced summer monsoon circulation. After &12.7 ka lake level decreased and a stagnant swampy environment was established. This is consistent with observations from Sumxi Co (Gasse et al., 1991; Van Campo and Gasse, 1993), where the Late Glacial is generally characterised by very low pollen production, sparse vegetation cover and high detrital input that is characteristic of dry climatic conditions. The return to an open, poorly evaporated lake took place at the beginning of the Holocene. This is consistent with results of Wang et al. (1990) from other western Tibetan lakes suggesting arid climate after 18.5 ka and a high lake level period between 16 and 13 ka. An older high lake level interval between '40}24 ka is reported from Aksayqin Hu, Bangong Co and Tianshuihai Hu (Li et al., 1989; Wang et al., 1990) and according to Rhodes et al. (1996) a #70 m lakeshore at Bangong Co could be dated C to 39 ka and Th/U to 41 ka, respectively. 5.2. Central Tibet The Atlas of Tibet Plateau (1990) shows many lakes in central Tibet that experienced higher lake levels in the past. Some of these higher lake shorelines should be threatened with caution as, for example, the reported 91 m higher lakeshore at Dagze Co (Fig. 2, No. 9) was identi"ed during "eld investigation as a geological fault line (Lehmkuhl, 1998c). The highest Pleistocene shoreline was situated at 62 m above modern lake and a #40 m level was tentatively dated to OIS 3 on basis of morphostratigraphy. Unfortunately, only a few dates for Late Pleistocene lake evolution in this area are available. Higher lake levels in the region of the central plateau may have occurred during the Late Pleistocene between 39}23 and 11}9 ka with lowest lake levels at about 22}16 ka (Wang, 1990). In the basin of lakes Peng Co and Pamu Co, a system of higher lakeshores can be identi"ed. The highest shoreline occurs at 87 m above recent lake level of Peng Co and 65 m of Pamu Co, respectively, when both lakes where connected together (HoK vermann and Lehmkuhl, 1994b). A 55}65 m higher shoreline at of Peng Co can be TL-dated by Frenzel (1994) to 48}45 ka. All lower shorelines show older ages between 68 ka (#18 m) and 142 ka (#12 m). The lakeshore sediments at the 12 m higher terrace at Yamzu Yum Co, however, gives ages between 22 and 20 ka. Chen (1986) dated a palaeosol situated on a 20.5 m higher shoreline of Nam Co to 14.5$4.6 ka. Above this soil horizon, lacustrine gravels accumulated suggesting a higher lake level after the soil developed, although the high standard deviation throws doubt on this age. Core results at Siling Co (Gue et al., 1993, cited by Gasse et al., 1996) suggest dry and cold climatic conditions between 12.2 and 10 ka. For the Dagze Co, a similar morphologi-

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cal sequences as in the Nam Co region can be derived from detailed "eld work (Lehmkuhl, 1998c) and will not be discussed in detail in this paper. 5.3. Northeastern Tibet The palaeoclimatic record and the lake level history of Qaidam Basin have been described above in terms of dust transport and accumulation mechanism. Interpreting the results of Liu (1991), the chronology of Late Pleistocene lake level changes at Qinghai Hu seems to be similar to that of the Qaidam Basin. The highest lakeshore of Qinghai Hu is situated about 141 m above the present lake level and is tentatively dated to &38 ka (Liu, 1991). Laminated lacustrine sediments and high pollen content indicate a higher lake level (#14 m) between 33.8 and 23.6 ka. Lake shrinking occurred after 23.6 to 14.8 ka when lakeshore sand and gravel accumulated. A dry and cold climate is deduced since 14.8 ka when loess accumulated widely in the lake area. At the southeastern end of the lake, yellow silt probably of aeolian origin accumulated approximately between 17 and 15 ka. According to Lister et al. (1991), a brackish lake and cold and windy climate existed until 11.8 ka. Studies on sediments show aeolian processes have been active at Qinghai Hu before 13.5 and until 12 ka. Since 12 ka lake level rose slightly, with some brief returns to drier intervals, until 6 ka (Lister et al., 1991; Kelts et al., 1989). A shoreline at 10 to 12 m above modern lake level at Xin Xin Hu and lacustrine sediments at Sogqen and Maligango have been TL-dated by Frenzel (1994) to 22 ka. A soil layer on the 10 to 12 m shoreline of the Xin Xin Hu has been dated to 17 ka with loess accumulation occurring at between 16 and 12 ka. This indicates lowering of lake level since 22 ka. Lehmkuhl et al. (1998) argued that sedimentation of loess started as early as 18 ka at the #23 m shoreline of this lake. 5.4. Inner Mongolia WuK nnemann et al. (1998) and WuK nnemann and Pachur (1998; Parts (9) and (10) of Fig. 9) published results from sediments, lakeshore features and cores from Gashun/Sogun Nur lake system located in Badain Jaran desert and Baijian Hu in Tengger desert, together with additional data from other locations. The chronology is based on 62 C dates. Highest lake levels of #28 to #32 m occurred between 41 and &33 ka in Gashun/Sogun Nur area and maximum extension of lake area at Baijian Hu between 40 and 23 ka. The lake level of Gaxun Nur was lowered 15 and 18 m at approximately 21 and 19 ka, respectively. Desiccation of Baijian Hu took place after 18 ka and at Gashun/Sogun Nur after 19 ka with dune sand accumulated continuously until 14 ka. In addition, no datable material has been found between 18.6 and 12.8 ka. This supports

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Fig. 10. Spatial distribution of lakes with evidences for higher lake levels during OIS 3. Ages are given in ka BP, indications of higher lake terraces in parentheses. Numbers according to Fig. 2; Q is Palaeo Qaidam lake.

hyperarid conditions, due to a lack of organic material. This hyperarid interval ended at Baijian Hu around 13 ka as the lake level rose again up to the #23 m level. A freshwater lake was established at Gashun/Sogun Nur after 11.3 ka. Sedimentological and palaeobotanical evidence in a sediment core shows that humid climatic conditions occurred at lake Manas between 37 and 32 ka (Rhodes et al., 1996). The age of the core is based on 9 C dates. The very low pollen content, episodically +1 accumulation of "ne-grained detritus of probably aeolian origin and the absence of datable material between 32 and 10 ka indicates generally dry climatic conditions during this period. The "rst humid impulse after longterm aridity occurred in Late Glacial times, approximately 12 ka (Rhodes et al., 1996). 5.5. Mongolia Naumann and Walther (2000) provided dates from geomorphological and sedimentological research at Bajan Nuur (Lake, Fig. 2 (27)). A #75}80 m lakeshore is dated to before 40 ka and lacustrine sediments with incorporated molluscs to 46.6 ka. Since 13.2 ka the Bajan Nuur lake began to "ll again. The highest Late Glacial level (#48 m) was dated by the radiocarbon method on

molluscs to 11.2 ka (`Postglacial Trangressiona). Similar results can be shown for the Uvs Nuur (Grunert et al., 2000). Tarasov et al. (1996) and Tarasov and Harrison (1998) describe lake level changes of Mongolian lakes (see Fig. 2). These are mainly attributed to the Holocene. However, higher lakeshores at Buir Nuur and Dood Nuur (39) and (31) of Fig. 2 are probably Middle Pleistocene in age. Lacustrine sediments and diatom assemblages at Hoton Nuur (25) of Fig. 2 provide evidence for a higher lake level between 12.2 ka and 10.1 ka, and 13.3 ka at Achit Nuur (26) of Fig. 2, respectively. In addition, sedimentological evidence from Buregkhanga area (1043E, 483N) show that aeolian activity was dominant between 40 and 30 ka, and colluvial activity from 30 ka until slightly younger than 24 ka (Feng et al., 1998). This suggests more humid conditions in the latter (30}24 ka) with respect to the earlier (40}30 ka) times.

6. Conclusions The di!erent palaeoclimatic evidence in Central Asia is summarised in Fig. 11. Glacier advances are observed for the OIS 4 and 2 (Benn and Owen, 1998) and for the Late Glacial around 15 ka. For the OIS 3 there were several

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dates on higher lake levels showing wetter climatic conditions (Fig. 10), but with a widespread range in this time frame. In some areas, high lake level persists until the OIS 2 (after 20 ka). However, there are regional di!erences, but generally all over Central Asia extremely arid climatic conditions occur in Late Glacial times, the socalled Kataglacial (c. 24}15 ka). For example, Owen et al. (1998) showed that permafrost developed in the Gobi of southern Mongolia in this period (22}15 ka, OSL dates) and permafrost degradation occurred during late Pleistocene time (13}10 ka). In Qaidam Basin cold and dry climatic conditions occur since 25 ka and hyperarid conditions can be assumed from 15}9 ka. There is also

139

some evidence that Lop Lake experienced a period of desiccation after deposition of lacustrine mud at 20.8 ka (Fang, 1991 and Fig. 9). The Qinghai Hu has a similar lake evolution as the Qaidam Basin; both have higher lake levels during OIS 3. They experienced arid conditions from 17}12 ka, similar to the lakes in the Gobi desert (Gashun/Sogun Nur), that dried out from 19}13 ka during which time 3 m of dune sand accumulated. Hofmann (1993) found evidence for more arid climate than today between 23 ka and soon after 10 ka at the Helan Shan (see (11) of Fig. 9). In addition, pollen and sedimentological evidence from western Loess Plateau indicates a smaller extent of the Gobi during cold and moist

Fig. 11. Landscape development including lake level changes, glacier #uctuations in the mountains (top) and basins (bottom) of Central Asia during the last 100,000 yr. The palaeosol sequences based on the Loess Plateau (Li et al., 1985) and the lacustrine phases a generalised form di!erent sources (see text for further details).

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periods 73}60 ka and 27}19 ka and maximum expansion of the Gobi desert 60}50 ka and 19}10 ka (Feng et al., 1998). TL dates from dune sand of sand-loess pro"les in the Mu Us Desert located north of the Loess Plateau prepared by Sun et al. (1998) indicate a hyperarid environment after 75$9 ka to before 55$4 ka and shortly before 17$1 until soon after 10$1 ka. C data +1 from the desert-loess transition zone of central Loess Plateau provide evidence for hyperarid climatic conditions before 13 ka due to sedimentation of 1.4 m of aeolian sand (Zhou et al., 1996). According to Fang (1991) extensive dune sands accumulated at the northern Qinghai-Tibetan Plateau region (Gonghe Basin) from 14.8 and 12.7 ka. In western Mongolia, results from the Bajan Nuur shows that the climate was dry before 13 ka (Naumann and Walther, 2000; Grunert et al., 2000). Observations on Late Quaternary alluvial fans in the Gobi of southern Mongolia (Owen et al., 1997b) suggest that more humid conditions may have occurred between 40 and 23 ka followed by a period of increased aridity. Synoptic interpretation of palaeoenvironmental information from southeastern Tibet given by Yan and Petit-Maire (1994) show that an arid episode probably occurred after LGM and persisted until around 14 ka. Frenzel (1994) suggested that moisture was availabile on the Tibetan Plateau but was lower during approximately 20 and 10 ka than today. However, in the western part of the Tibetan Plateau there is local evidence for a return to wetter conditions since approximately 16}15 ka. In the centre part of the Plateau there are a few dates on higher lake levels, but there is still a lack of serious investigation in this vast area. This change towards more humid and warmer conditions began in the eastern part of the Plateau around 13}10 ka with a gradual reinforcement of the summer monsoon in central China (e.g. Pachur et al., 1995). Rapid, short-termed climatic oscillations such as the Younger Dryas event may be recorded in the loess pro"les (An et al., 1993; Zhou et al., 1996) or in lacustrine sediments in western Tibet (e.g. Gasse et al., 1991). The comparison of dust #ux changes, palaeoprecipitation values and landscape evolution at Qaidam Basin shows that the interpretation of increasing dust #ux due to increasing aridity in the dust source regions is not conclusive. First of all, Qaidam Basin is characterised by hyperarid climatic conditions today, as well as the Taklimakan and Gobi deserts. The increase in dust #ux at the Dunde icecap started as early as 30 ka, when climate was more humid than today. This increase is attributed to expansion or even the establishment of dust producing environments. The peaks of dust #ux correlate with the transition from colder and wetter to drier climatic conditions, which include transition from high lake levels (storage of "ne grained clastics) to low lake levels (de#ation of formerly stored silt material). Depending on the di!erent types of climatic induced landscapes, de#ated material

could be compensated to some degree. The decrease of dust #ux values is attributed mainly to long-term aridity and associated landscape types that are unable to compensate formerly de#ated material. The widespread distribution of gravel in the Gobi of Central Asia originated mainly from alluvial fans and fanglomerates (pediment region). As shown above a transition from wetter (mainly OIS 3 and beginning OIS 2"Anaglacial period) to arid climatic conditions (OIS 2, Kataglacial period) occurred in most areas of Central Asia. It is possible that the local dust #ux model from the Qaidam Basin (Fig. 7) can be assigned for all over the area with slight modi"cations within this time frame. In addition, in most of the marine deep-sea records, and ice core drillings, even in Antarctica (Vostok), the highest dust #ux values take place during the cold stages of the Last Glacial (OIS 4 and 2). The Anaglacial time (c. 32}24 ka) shows wetter conditions, that could have been induced by an increase of the westerlies and a weakening of the winter monsoon. This "rst in#uenced the western part of Central Asia; the latter in#uenced the mainland of China and the eastern part of Central Asia. During Kataglacial times (c. 24}15 ka) a strengthening of the winter monsoon due to a signi"cant displacement of the Siberia high-pressure to the south could have caused the dry climatic conditions in north and central China (An et al., 1991b; Ding et al., 1995; Pachur et al., 1995; Rost, 1998). The dominating N to NW winter monsoon winds were intensely increased and Central Asia and northern China was controlled by a cold-dry and windy climate, mainly during springtime. The apparent increase in aridity was at least partly enhanced by the eastward movement of the coastline of the Yellow Sea during this period. This resulted in the intensi"ed dust transport from the desert areas of Central Asia and the increased accumulation of the younger (Malan) Loess sequences in China. The di!erent variations, especially in precipitation, might have been dominated by the dynamic and thermal e!ects of the Siberia High. This could provide a signi"cant increase in the gradient of the wind and temperature "elds, with the high-pressure cell becoming more persistant in the winter and springtime. During more humid periods in Central Asia this pressure system might be relatively weak and farther north, enhancing cyclonic precipitation by the westerlies or/and the monsoon in springtime. Holocene palaeosols developed as aeolian loess continued to accumulate, although at a reduced rate and the climate might have been wetter than today due to an intensi"ed summer monsoon. Three major intervals of soil formation occurred in c. 9900}8000 BP, 7400}4600 BP, and 3400}2000 BP (Liu et al., 1985; see Fig. 11). These can be compared with lake level #uctuations. The latter period might correlate with a Late Holocene cooling period, the so-called Neoglacial from 4000 to 3000

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BP. This time is determined from glacier advances, pollen records, and intensi"ed soli#uction in the mountains (e.g. Pu, 1991; Wang and Fan, 1987; Sun and Chen, 1991; Lehmkuhl, 1995).

Acknowledgements The authors wishes to thank the Chinese Academy of Sciences (CAS), the Deutsche Forschungsgemeinschaft (DFG), the Gesellschaft fuK r Technische Zusammenarbeit (GTZ) and the Max-Planck-Gesellschaft (MPG) for their "nancial support of the expeditions. The compilation of the references on palaeoclimate of Central Asia was part of a BMBF-Project on `terrestrial Palaeoclimatea in 1995}1997 (Prof. Dr. J. HoK vermann; FoK rdernummer: 07VKV/01A, Unterauftrag 21178.4). In addition, we would like to thank also Prof. Dr. M.A. Geyh (Hannover) for the radiocarbon dating and discussions. The TLsamples (TGD) were analysed by Mrs. Lu Liangcai and Prof. Huang Baolin; Institute of Geochemistry, Chinese Academy of Sciences, Guanzhou. We should like to thank Prof. Dr. L.A. Owen for helpful suggestions in improving the English and valuable comments on the scienti"c content.

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