Holocene climate change inferred from stratigraphy and OSL chronology of aeolian sediments in the Qaidam Basin, northeastern Qinghai–Tibetan Plateau

YQRES-03481; No. of pages: 12; 4C: Quaternary Research xxx (2013) xxx–xxx

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Holocene climate change inferred from stratigraphy and OSL chronology of aeolian sediments in the Qaidam Basin, northeastern Qinghai–Tibetan Plateau LuPeng Yu a,b,c, ZhongPing Lai a,c,d,⁎ a

CAS Key Laboratory of Salt Lake Resources and Chemistry, Qinghai Institute of Salt Lakes, Chinese Academy of Sciences, Xining 81000, China Qinghai Geological Survey Institute, Xining 810012, China State Key Laboratory of Cryospheric Sciences, Cold and Arid Regions Environmental and Engineering Research Institute, Chinese Academy of Sciences, Lanzhou 730000, China d State Key Laboratory of Loess and Quaternary Geology, Institute of Earth Environment, Chinese Academy of Sciences, Xi'an 710075, China b c

a r t i c l e

i n f o

Available online xxxx Keywords: OSL dating Aeolian activity Holocene climate change Asian summer monsoon Qaidam Basin Qinghai–Tibetan Plateau

a b s t r a c t Paleoclimatic reconstruction based on aeolian sediments in the eastern Qaidam Basin (QB) has been hindered by the limited chronological data. Here we present 61 Optically Stimulated Luminescence (OSL) ages. On the basis of these OSL ages and the lithologic stratigraphy, we propose the ‘effective moisture index (EMI)’ for aeolian sediments to reconstruct the effective moisture change. Based on the EMI from twelve sections, the effective moisture change, moisture sources and relevant mechanisms for paleoclimatic change in the eastern QB are discussed. The results indicate that (1) aeolian deposition started at least before 12.4 ± 0.7 ka during the deglaciation, the paleosols developed at the early and mid-Holocene, and aeolian sand and loess accumulated at mid- and late Holocene; (2) effective moisture history was: hyper-arid at 12.8–11.6ka, humid and variable at 11.6–8.3 ka, moderately humid and stable at 8.3–3.5 ka, and increasingly arid at 3.5–0 ka; (3) the effective moisture change was mainly controlled by the Asian summer monsoon (ASM), which mainly followed the change of Northern Hemispheric summer insolation, and the westerlies strengthened and increased the aridity in the QB when the ASM shrank. © 2013 University of Washington. Published by Elsevier Inc. All rights reserved.

Introduction The Qaidam Basin (QB) in the northeastern Qinghai–Tibetan Plateau (QTP) (Figs. 1A and B), which encompasses one of the highest (elevation N2800 m) and driest (b50 mm average annual precipitation in its western part) deserts on the earth, is a major source area for modern Asian mineral dust (Chen et al., 2007). Clastics from the QB not only served as a major dust source for the Chinese Loess Plateau (Kapp et al., 2011; Pullen et al., 2011; An et al., 2012a; Lai et al., 2013), but also lead to accumulation of the widely distributed aeolian sediments in the eastern QB. Located at a triple junction of influences from the East Asian summer monsoon (EASM), the westerlies, and the Indian summer monsoon (ISM) (Fig. 1A) in summer, the QB is sensitive to climatic variability. Though most part of the QB is largely influenced by the westerlies (Chen and Bowler, 1985; Sun et al., 1998; Zhao et al., 2007), its eastern part, located near the modern Asian summer monsoon (ASM, including EASM and ISM) limit (Figs. 1A and B), might be easily influenced by the enhanced ASM during the Holocene. Previous studies proposed ⁎ Corresponding author at: CAS Key Laboratory of Salt Lake Resources and Chemistry, Qinghai Institute of Salt Lakes, Chinese Academy of Sciences, Xining 81000, China. E-mail address: [email protected] (Z.P. Lai).

that even the enhanced ASM during the warm and humid period of Holocene could not penetrate into the interior of the western desert regions in China, and no Holocene soils were found in the Qaidam deserts (Gao et al., 1996; Sun et al., 1998). However, we have found and dated widely distributed Holocene paleosols around the Tiekui desert (Figs. 1C and 2D) in the eastern QB in this study. The study of paleoclimate in the QB has attracted international attention since the 1980s (e.g., Chen and Bowler, 1985; Bowler et al., 1986; Chen and Bowler, 1986; Zhao et al., 2007; Sun et al., 2010; Lai et al., 2013), with most studies focused on lake sediments, however, the chronology of lacustrine sediments might subject to the hardwater effect of radiocarbon dating (Yang et al., 2011). Although aeolian records might not be as continuous as those of lake sediments, they can be directly dated using Optically Stimulated Luminescence (OSL) techniques and their connection with climatic change is relative straightforward (Lu et al., 2011). However, due to limited age control, aeolian processes and their relevance to paleoclimatic change in the QB are yet poorly understood. Hao et al. (1998) reported one 14C age (1.465 ± 0.060 cal ka BP at 70–80 cm) and two thermoluminescence (TL) ages (32.0 ka at 250–260 cm and 50.2 ka at 340–350 cm), and also inferred that the paleosols at 430–530 cm was developed during the last interglaciation in the eastern QB. However, such a low accumulation rate

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Please cite this article as: Yu, L.P., Lai, Z.P., Holocene climate change inferred from stratigraphy and OSL chronology of aeolian sediments in the Qaidam Basin, northeastern Qinghai–Tibetan Plateau, Quaternary Research (2013), http://dx.doi.org/10.1016/j.yqres.2013.09.006

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Figure 1. (A), map showing the location of the Qinghai–Tibetan Plateau (QTP), the Qaidam Basin (QB), the Chinese Loess Plateau (CLP) and the dominant circulation systems of the westerlies (W), the Indian summer monsoon (ISM), the east Asian summer monsoon (EASM) and the east Asia winter monsoon (EAWM). The yellow square denotes the location of the QB, the red dotted line marks the modern Asian summer monsoon limit (modified from Gao, 1962). (B), the DEM map reveals the geomorphology of the QB and the lakes mentioned in the text (1: Da Qaidam Lake, 2: Xiao Qaidam Lake, 3: Hurleg Lake, 4: Gahai Lake, 5: Bieletan Lake, 6: Chaka Lake, and 7: Qinghai Lake). The yellow square indicates the study area of Tiekui Desert shown in (C), which displays the positions of the profiles in this study (XRH1–XRH2 (98°10′29″ E, 36°25′50″ N, 3174 m), XRH3 (98°28′16″ E, 36°26′12″ N, 3506 m), XRH4 (98°27′45″ E, 36°26′22″ N, 3494 m), XRH5 (98°12′21″ E, 36°26′14″ N, 3201 m) and SYK1 (97°54′00″ E, 36°21′38″ N, 3060 m), in Yu and Lai (2012) (WLS1 (97°46′13″ E, 36°01′11″ N, 3105 m), XXT2–XXT2-2 (98°03′35″ E, 36°15′53″ N, 3237 m), BDB1–BDB1-2 (98°00′23″ E, 36°02′34″ N, 3202 m), and an unpublished section (HBC1 (98°07′54″ E, 36°25′41″ N, 3143 m). (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)

(ca. 40–70 mm/ka) of aeolian deposit at the margin of this hyper-arid and windy basin seems unreasonable according to our former study in this region (Yu and Lai, 2012). Zeng et al. (1999, 2003) obtained four TL ages (20.20 ± 4.04, 18.01 ± 3.60, 14.87 ± 2.97 and 11.36 ± 2.27 ka) from the dune sands in the eastern QB and pointed out that the dune sands were deposited during the Last Glacial Maximum (LGM) and Younger Dryas (YD) event. Zeng (2006) dated an archeological site in Nuomuhong, southern QB, and his two 14C ages (3.229 ± 0.082 and 3.477 ± 0.063 cal ka BP) of the two ash layers indirectly denoted the ages of aeolian sand in that section. Owen et al. (2006) reported an OSL age of ~15ka from an aeolian sand wedge at Lenghu in the northern QB. Loess on the Golmud River terraces were reported to be formed since 8 ka (Owen et al., 2006) and 13.9 ka (Chen et al., 2011) based on OSL dating in the Kunlun Mountains to the south of the QB. Niu et al. (2010) reported four OSL and four 14C ages from paleosols sections in the eastern QB, and these OSL ages were in agreement with the 14C ages for ages younger than 5 ka. Unfortunately, no dating details were provided for all these OSL ages. Zhou et al. (2012) revealed that the linear dunes in the central QB formed at 3.21–0.81 ka by OSL dating. Yu et al. (2013) dated paleo-dunes in the middle and southwestern margin of the QB, and OSL ages displayed those paleo-dunes accumulation started at ca. 3–4 ka, in response to the arid climate during the late Holocene. Yu and Lai (2012) provided 28 OSL ages from eight aeolian

sand–loess sections in the eastern QB, and their results showed that (1) the aeolian sediment accumulation started during the deglaciation; (2) aeolian sand was accumulated at 12.4–11.5 and 10–8 ka; (3) loess accumulation started at 10 ka and lasted to 0.45 ka at least, while large-scale loess accumulation occurred at ~8–4.5 ka. However, the reconstructed humidity during 10–8 ka seemed ambiguous due to the limited aeolian record. Previous works were mainly focusing on loess and aeolian sand, while the paleosols were mostly ignored or not discovered. Development of the paleosols in the hyper-arid QB requires higher effective moisture; as a result, chronology of the paleosols is of great significance to the reconstruction of paleoclimate, and to discuss the source of the increased moisture. Two atmospheric circulation systems, i.e., the mid-latitude westerlies and the ASM, play key roles in Northern-Hemisphere climatic change. However, their influence on the climate in the QB still remains unclear. Though a number of studies have suggested that the warm and humid period during the Holocene was time-transgressive across the monsoon regions (e.g., An et al., 2000; Hong et al., 2003, 2005; Wang et al., 2010), high-resolution and well-dated speleothem records from both Oman (Fleitmann et al., 2003) and China (Dykoski et al., 2005; Wang et al., 2005; Shao et al., 2006; Hu et al., 2008; Wang et al., 2008) have displayed a general drying trend during the Holocene with

Please cite this article as: Yu, L.P., Lai, Z.P., Holocene climate change inferred from stratigraphy and OSL chronology of aeolian sediments in the Qaidam Basin, northeastern Qinghai–Tibetan Plateau, Quaternary Research (2013), http://dx.doi.org/10.1016/j.yqres.2013.09.006

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Figure 2. (A) Wind-eroded yardangs in the western Qaidam Basin. (B) and (C) show the aeolian sections and modern Tiekui Desert in the eastern Qaidam Basin, respectively. (D) displays the stratigraphy of aeolian sediments in XRH2 section in the eastern Qaidam Basin.

a humid early and mid-Holocene and an arid late Holocene, in the regions influenced by the ASM. Whereas in westerlies-controlled arid Central Asia, the climate was dry during the early Holocene, wetter (less dry) during the mid-Holocene and moderately wet during the late Holocene (Chen et al., 2008). Consequently, the distinct characteristic of the climate patterns between monsoonal Asia and Central Asia is the difference in the humidity during the early Holocene. Comparing the effective moisture change displayed by the aeolian sediments in this region with that of the monsoonal Asia and Central Asia can help to evaluate the dominating climate system during the Holocene in the QB. Conflicts still exist about Holocene climatic change in the QB. Based on 14C dating, Zhao et al. (2007) showed that climate was dryer during the early Holocene and wetter during the Late Holocene in the Hurleg Lake (3 in Fig. 1B), while other studies indicated a wetter early and mid-Holocene in the Chaka Lake (6 in Fig. 1B, Liu et al., 2008), Gahai Lake (4 in Fig. 1B, Chen, 2010; Li, 2011), Bieletan Lake and Da Qaidam Lake (5 and 1 in Fig. 1B, Huang et al., 1981). Here, we present stratigraphic analysis and detailed OSL chronology in order to understand the evolution of the aeolian sediments, and to reconstruct the effective moisture history for the QB. Study area The QB, a large intermontane depression located at the northern margin of the Qinghai–Tibetan Plateau (QTP) (Fig. 1A), has an average elevation of 2800 m asl at the basin floor, and is surrounded by the over 5000 m Qilian, Kunlun and Altun Mountains (Fig. 1B). The basin covers an area of 1.2 × 105 km2, and has a catchment of about 2.5 × 105 km2. The annual mean precipitation and evaporation of the hyper-arid central basin are about 26 mm and 3000–3200 mm, respectively. The landforms are mainly composed of yardangs (Fig. 2A), saltlake playas and aeolian deposits (Figs. 2B and C) from west to east. The westerlies are the dominant atmospheric circulation over the basin year round, with prevailing strong north and/or northwesterly winds resulting from the invasion of cold air masses from highlatitude regions (Wu et al., 1985), and nearly all the dust storms occur

in spring (the frequency of dust storms weather is ~40 day/yr) (Chorographic committee of Dulan County, 2001). Severe wind erosion is demonstrated in the western QB by the presence of extensive fields of yardangs in the west part of the basin (Fig. 2A) and the surface age of ~100 ka in the Chaerhan salt lake which is the deposition center of the basin in the Quaternary (Lai et al., 2013). Clastics eroded from these lacustrine strata are the major source of modern dust storms, and might be a dominant source of the sediments in the Chinese Loess Plateau, which is located at downwind of the QB (Kapp et al., 2011; Pullen et al., 2011; An et al., 2012a; Lai et al., 2013). A portion of this winderoded material was redistributed to, and stored within the eastern QB (Kapp et al., 2011; Yu and Lai, 2012), where the Tiekui Desert is located (Figs. 1B and C). The Tiekui Desert, the largest desert in the QB, is at 2800–3300m asl, with playas below 2800 m to the west and mountains over 4000 m to the east (Fig. 1B). Mobile dunes are widely distributed inside the desert (Fig. 2C), and the 1–2 m thick loess/paleosol deposits are primarily distributed around the desert margin, with aeolian sand layers distributed under or intercalated within the loess/paleosol (Yu and Lai, 2012) (Figs. 2B and D). In the eastern margin of the Tiekui Desert (Fig. 1C), the average annual precipitation of the Xiariha Town is 240.8 mm, and 52–80% of the annual precipitation is received in summer (Chorographic committee of Dulan County, 2001). Sections and samples Detailed surveys were conducted in the Tiekui Desert, especially in the eastern desert margin, where our profiles are located (Fig. 1C). Identification of the paleosol, weak expressed soil, loess, and aeolian sand in the field was based on the grain size, structure, color, water content, and compactness (Fig. 2D). The paleosol is usually rufous and bronze in color, and with finer grain size, higher water content, and compact and with multihole soil structure. The weakly expressed soil has the soli structure but has lower organic content, coarser grain size, and lighter color compared with the paleosol. The loess is usually gray or yellow in color, with many vertical fissures and weaker soil structure, and without horizontal beddings. The aeolian sand is incompact, coarse, and

Please cite this article as: Yu, L.P., Lai, Z.P., Holocene climate change inferred from stratigraphy and OSL chronology of aeolian sediments in the Qaidam Basin, northeastern Qinghai–Tibetan Plateau, Quaternary Research (2013), http://dx.doi.org/10.1016/j.yqres.2013.09.006

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sometimes with inclined beddings, therefore is easy to be identified. The mean grain size, color (lightness, redness and yellowness) and magnetic susceptibility (MS) were analyzed in the laboratory to confirm the field identification. Comparison between stratigraphic identification in the field and measured proxies in the laboratory (including mean grain size, MS and color) was conducted on the lower part of XRH1 section, and the results (Fig. 3) confirm our field identification. Variation in the mean grain size records in other sections was also consistent with the field observations (Fig. 4). Six sections consisting of paleosols, weakly expressed soil, loess and aeolian sand (see Figs. 1C, 2D and 4) were selected and 61 OSL samples were collected in order to provide a systematic chronology (Fig. 4). The 750 cm thick Shangyakou1 (SYK1) section, located in the modern Tiekui Desert, is mainly composed of aeolian sands, and intercalated with loess (at 140–230 cm), weakly expressed soil (at 285–330 and 363–397 cm) and paleosols (most of the part at 450–605 cm). 15 OSL samples were collected from this section. Another OSL (SYK1-0) sample was taken from a modern dune at the depth of 30 cm close to this section. The Xiariha (XRH1, 2, 3, 4 and 5) sections are located on the fluvial sediments of the Xiariha river terraces to the east of the Tiekui Desert, where paleosols of the QB mainly are distributed. The XRH sections consist of paleosols and weakly expressed soil at the lower part and aeolian sand/loess at the upper part (Fig. 4), which suggest a general drying trend during the period of accumulation. These paleosols are usually intercalated with aeolian or fluvial sands. The XRH1 and XRH 2 sections are from the same site, but the XRH2 section locates on a higher terrace.

We did not reach the base of the aeolian sediments in the XRH4 and XRH 5 sections. 46 OSL samples were collected from the XRH sections. A total of 57 OSL samples were taken from the aeolian sediments (11 samples from loess, 20 samples from aeolian sand, and 26 samples from paleosols). Four samples were taken from the underlying fluvial sediments, which could provide age estimates for both the terrace and the maximum age of the onset of the overlying aeolian sand. The detailed positions for all OSL samples are shown in Figure 4. All luminescence samples were collected by hammering steel tubes (~22 cm long cylinder with a diameter of ~5 cm) into freshly cleaned vertical sections. The tubes were then wrapped to avoid light exposure. Bulk samples were also collected in each location for water content and dose rate analysis. OSL dating OSL sample preparation and measurement techniques In the laboratory of Qinghai Institute of Salt Lakes, Chinese Academy of Sciences, the unexposed middle part of the tube was used to extract minerals for equivalent dose (De) determination. The samples were treated with 10% HCl and 30% H2O2 to remove carbonates and organics, respectively. Grain-size fraction of 38–63 or 90–125 μm was extracted by wet sieving. The fraction of 38–63 μm was etched by 35% H2SiF6 for about two weeks to remove feldspars (Lai and Wintle, 2006; Lai et al., 2007a; Roberts, 2007). The fraction of 90–125 μm was treated with 40% HF for 45 min to remove feldspars and the alpha-irradiated outer layer (~10 μm). The resulting quartz grains were washed with 10% HCl to remove fluoride precipitates. The purity of quartz grains was checked by infrared (830 nm) stimulation, and any samples with obvious infrared stimulated luminescence (IRSL) signals were retreated with H2SiF6 again to avoid De underestimation (Lai and Brückner, 2008). For the modern sample SYK0, the OSL signals from quartz were too low to be detected, so this sample was tested by IRSL, and the fraction of 38–63 μm without H2SiF6 treatment was used. The pretreated grains were then mounted on the center part (~0.5 cm diameter) of stainless steel disks (1 cm diameter) using silicone oil. OSL measurements were made using an automated Risø TL/OSL-DA20 reader equipped with blue diodes (λ = 470 ± 20 nm) and IR laser diodes (λ = 830 nm). The luminescence was stimulated by blue LEDs at 130°C for 40s, and detected using a 7.5mm thick U-340 filter (detection window 275–390 nm) in front of the photomultiplier tube. Ninety percent diode power was used. Irradiations were carried out using a 90Sr/ 90 Y beta source in the Risø reader. A preheat plateau test was conducted on sample XRH1-G, and a preheat plateau was clearly identified from 240 to 260 °C. Preheat was at 260 °C for 10 s for natural and regenerative doses, and cut-heat was at 220 °C for 10 s for test doses. Signals of the first 0.64 s stimulation were integrated for growth curve construction after background subtraction (last 10 s). The concentrations of U, Th and K were measured by neutron activation analysis. For the 36–63 μm grains, the alpha efficiency value was taken as 0.035 ± 0.003 (Lai et al., 2008). The cosmic-ray dose rate was estimated for each sample as a function of depth, altitude and geomagnetic latitude (Prescott and Hutton, 1994). The dose rates are shown in Table 1. Equivalent dose determination

Figure 3. Comparison of stratigraphic identification in the field with measured proxies in the lab (including Mean Grain Size, magnetic susceptibility (MS) and color: yellowness, redness and lightness). (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)

In the current study, the combination of SAR protocol (Murray and Wintle, 2000) and the Standard Growth Curve (SGC) method (Roberts and Duller, 2004; Lai, 2006; Lai et al., 2007b; Yu and Lai, 2012), named as SAR–SGC method (Lai and Ou, 2013), was employed for De determination. In this method, for each sample, 4–6 aliquots were measured using the SAR protocol to get 4–6 growth curves which were then averaged to construct a SGC for this individual sample. Then, 8–16 more

Please cite this article as: Yu, L.P., Lai, Z.P., Holocene climate change inferred from stratigraphy and OSL chronology of aeolian sediments in the Qaidam Basin, northeastern Qinghai–Tibetan Plateau, Quaternary Research (2013), http://dx.doi.org/10.1016/j.yqres.2013.09.006

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Figure 4. Stratigraphy, OSL ages and Mean Grain Size of the sections used for effective moisture index reconstruction. The WLS1, XXT2, XXT2-2, BDB1 and BDB1-2 sections are from Yu and Lai (2012), and the HBC1 section (unpublished data). The ages marked with underlines are not consistent with the stratigraphic order, and the ages marked with * mean measured with 90–125 μm fraction.

aliquots were measured to obtain the values of test-dose corrected natural signals only, and each of the values could be matched in the SGC to obtain a De. De results determined by the SGC are well in agreement with those by the SAR protocol within 10% for most samples (Fig. 5), suggesting that the SAR–SGC method could be used for De determination for samples from the QB. For all samples, the final De is the mean of all SAR Des and SGC Des. The validity of the SAR protocol was tested with a ‘dose recovery test’ (Murray and Wintle, 2003) on sample XRH1-D, XRH1-G and XRH5-12, and six aliquots of each sample were tested. The given laboratory dose were 32, 15 and 25 Gy, and the measured De were 31.9 ± 1.74, 14.4 ± 0.80 and 23.7 ± 1.77 Gy, respectively. Thus, the ratios of the measured to the given doses were 0.997 ± 0.054, 0.960 ± 0.053 and 0.948 ± 0.071, suggesting that the SAR protocol is suitable for De determination. Figure 6A and C show typical OSL decay curves of XRH4-2 and XRH48, respectively, which indicate that the OSL signals were from the fast component. Figure 6B and D show the SGCs for sample XRH4-2 and XRH4-8, respectively, and all the growth curves and SGCs of the other samples can be also well fitted using the exponential plus linear function. Recuperation was calculated by comparing the sensitivity-

corrected OSL signal of 0 Gy to the sensitivity-corrected natural signal to check the thermo-transferred signals. Finally the recuperation was b3% for all samples, which was negligible. The ‘recycling ratio’ was introduced to check for sensitivity change correction (Murray and Wintle, 2000), and for most aliquots, the recycling ratios fall into the acceptance range of 0.9–1.1. Dating results and aeolian activities since the late deglaciation Dating results OSL dating results are listed in Table 1, and are also shown in Figure 4. The coherence of OSL ages from 38–63μm and 90–125μm fractions of the two samples (XRH1-B and XRH2-A, Fig. 4 and Table 1) demonstrate that both of the two fractions are suitable for dating the aeolian sediments in the QB. Ages of the aeolian sediments fall within the range of 12.1 ± 0.7–0.023 ± 0.002 ka, while ages of the fluvial sediments fall within the range of 18.2 ± 1.0–9.4 ± 0.7 ka. The age of 0.023 ± 0.002 ka at the depth of 30 cm in a modern dune demonstrates that the aeolian sediments were well bleached before deposition and

Please cite this article as: Yu, L.P., Lai, Z.P., Holocene climate change inferred from stratigraphy and OSL chronology of aeolian sediments in the Qaidam Basin, northeastern Qinghai–Tibetan Plateau, Quaternary Research (2013), http://dx.doi.org/10.1016/j.yqres.2013.09.006

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Table 1 Environmental radioactivity and OSL dating results. Sample ID

Depth (m)

K (%)

Th (ppm)

U (ppm)

Water content (%)

Dose rate (Gy/ka)

Aliquot number

De (Gy)

OSL age (ka)

SYK1-0IRSL SYK1-1 SYK1-2 SYK1-3 SYK1-4 SYK1-5 SYK1-6 SYK1-7 SYK1-8 SYK1-A* SYK1-B SYK1-B* SYK1-C SYK1-D SYK1-E* SYK1-F XRH1-1 XRH1-E XRH1-F XRH1-G XRH1-H XRH1-2 XRH1-A XRH1-B XRH1-C XRH1-3 XRH1-D XRH2-A XRH2-A* XRH2-1 XRH2-B XRH2-C XRH2-D XRH2-E XRH3-1 XRH3-2 XRH3-3 XRH3-4 XRH3-5 XRH4-1 XRH4-2 XRH4-3 XRH4-4 XRH4-5 XRH4-6 XRH4-7 XRH4-8 XRH5-1 XRH5-2 XRH5-3 XRH5-4 XRH5-5 XRH5-6 XRH5-7 XRH5-8 XRH5-A XRH5-9 XRH5-B XRH5-10 XRH5-11 XRH5-C XRH5-12 XRH5-D

0.30 0.35 1.00 1.70 2.05 2.60 3.40 4.20 4.90 7.50 6.40 6.40 5.90 5.40 4.90 1.5 0.50 0.80 1.20 1.60 2.15 2.50 2.95 3.60 3.90 4.45 4.75 0.60 0.60 1.18 1.60 2.15 2.75 3.35 0.55 1.05 1.65 2.70 3.30 0.45 1.10 2.50 3.00 3.65 4.25 3.30 5.90 0.50 0.95 1.30 1.70 2.75 3.20 3.45 4.00 4.35 4.65 5.10 5.50 5.80 6.15 6.45 7.00

2.04 ± 0.07 1.79 ± 0.08 1.75 ± 0.08 1.68 ± 0.07 1.66 ± 0.07 1.64 ± 0.08 1.61 ± 0.06 1.62 ± 0.07 1.79 ± 0.07 1.67 ± 0.06 1.90 ± 0.07 1.90 ± 0.07 2.10 ± 0.07 1.93 ± 0.07 1.76 ± 0.07 1.81 ± 0.07 1.69 ± 0.07 1.59 ± 0.09 2.15 ± 0.10 1.69 ± 0.09 1.82 ± 0.09 1.58 ± 0.07 1.86 ± 0.09 2.15 ± 0.11 1.88 ± 0.10 1.34 ± 0.06 1.42 ± 0.09 1.78 ± 0.06 1.78 ± 0.06 1.52 ± 0.07 1.71 ± 0.06 1.66 ± 0.06 1.74 ± 0.06 1.71 ± 0.06 1.63 ± 0.09 1.54 ± 0.08 1.66 ± 0.09 1.60 ± 0.09 1.59 ± 0.09 1.51 ± 0.08 1.56 ± 0.08 1.51 ± 0.08 1.65 ± 0.08 1.59 ± 0.08 1.63 ± 0.08 1.43 ± 0.08 1.56 ± 0.08 1.63 ± 0.09 1.80 ± 0.09 1.91 ± 0.10 1.86 ± 0.09 2.12 ± 0.10 1.94 ± 0.10 1.84 ± 0.09 1.71 ± 0.09 1.43 ± 0.06 1.65 ± 0.09 1.44 ± 0.05 1.91 ± 0.09 1.82 ± 0.09 1.44 ± 0.05 1.81 ± 0.09 1.86 ± 0.06

4.83 ± 0.16 5.49 ± 0.20 5.87 ± 0.22 9.51 ± 0.29 9.31 ± 0.27 7.86 ± 0.24 6.98 ± 0.21 6.65 ± 0.22 8.91 ± 0.28 9.89 ± 0.25 5.83 ± 0.18 5.83 ± 0.18 10.18 ± 0.26 8.48 ± 0.22 6.14 ± 0.18 9.48 ± 0.25 9.61 ± 0.29 8.55 ± 0.48 12.37 ± 0.59 9.92 ± 0.52 9.68 ± 0.50 9.62 ± 0.28 10.14 ± 0.51 12.27 ± 0.60 12.08 ± 0.62 10.67 ± 0.32 10.40 ± 0.63 10.53 ± 0.27 10.53 ± 0.27 7.95 ± 0.29 9.88 ± 0.26 9.48 ± 0.26 8.97 ± 0.24 7.83 ± 0.22 11.03 ± 0.62 10.41 ± 0.56 11.06 ± 0.59 9.30 ± 0.51 7.21 ± 0.39 10.29 ± 0.52 10.31 ± 0.53 9.10 ± 0.47 9.77 ± 0.49 10.59 ± 0.53 10.95 ± 0.56 9.26 ± 0.51 9.83 ± 0.53 9.05 ± 0.49 8.12 ± 0.42 12.24 ± 0.61 11.61 ± 0.58 11.56 ± 0.54 11.27 ± 0.55 11.26 ± 0.56 12.40 ± 0.64 10.25 ± 0.28 12.39 ± 0.64 9.94 ± 0.25 12.03 ± 0.59 10.71 ± 0.52 9.94 ± 0.25 11.57 ± 0.58 12.14 ± 0.30

1.38 ± 0.12 1.50 ± 0.20 1.63 ± 0.19 3.06 ± 0.23 3.04 ± 0.24 2.13 ± 0.21 1.83 ± 0.20 1.72 ± 0.18 2.71 ± 0.23 1.94 ± 0.13 1.69 ± 0.12 1.69 ± 0.12 3.18 ± 0.16 2.56 ± 0.14 1.67 ± 0.11 2.67 ± 0.13 3.15 ± 0.25 2.41 ± 0.21 4.50 ± 0.27 4.04 ± 0.26 3.60 ± 0.24 2.38 ± 0.22 3.07 ± 0.22 4.06 ± 0.25 3.79 ± 0.26 2.73 ± 0.22 2.45 ± 0.20 2.98 ± 0.14 2.98 ± 0.14 2.03 ± 0.24 2.85 ± 0.15 2.95 ± 0.15 2.95 ± 0.14 1.98 ± 0.13 2.49 ± 0.20 3.05 ± 0.23 2.95 ± 0.23 2.45 ± 0.21 2.25 ± 0.20 3.32 ± 0.22 3.50 ± 0.21 2.96 ± 0.21 3.57 ± 0.22 3.29 ± 0.20 3.11 ± 0.20 2.64 ± 0.19 2.67 ± 0.19 2.47 ± 0.20 2.64 ± 0.22 3.13 ± 0.23 3.43 ± 0.24 3.60 ± 0.24 3.47 ± 0.24 3.21 ± 0.24 3.21 ± 0.24 2.59 ± 0.15 3.02 ± 0.21 2.67 ± 0.14 3.49 ± 0.24 3.05 ± 0.22 2.67 ± 0.14 3.17 ± 0.23 3.07 ± 0.17

5±2 5±2 5±2 5±2 5±2 5±2 5±2 5±2 10 ± 5 10 ± 5 5±2 5±2 5±2 5±2 5±2 8±3 10 ± 5 5±2 5±2 5±2 5±2 5±2 12 ± 5 17 ± 5 16 ± 5 4±2 5±2 2±1 2±1 3±1 10 ± 5 9±5 12 ± 5 3±2 3±2 7±2 6±2 7±2 3±1 6±2 7±2 11 ± 5 14 ± 5 9±5 11 ± 5 4±2 9±5 6±2 4±2 6±2 7±2 9±4 7±2 7±2 7±2 5±2 14 ± 5 10 ± 5 15 ± 5 15 ± 5 7±2 22 ± 8 24 ± 8

3.25 ± 0.17 2.88 ± 0.16 2.87 ± 0.16 3.44 ± 0.17 3.39 ± 0.17 3.05 ± 0.17 2.88 ± 0.16 2.80 ± 0.15 3.15 ± 0.23 2.84 ± 0.12 3.10 ± 0.16 2.80 ± 0.13 3.44 ± 0.25 2.91 ± 0.21 2.70 ± 0.12 3.38 ± 0.18 3.56 ± 0.18 3.19 ± 0.18 4.08 ± 0.29 3.65 ± 0.19 3.55 ± 0.26 3.23 ± 0.15 3.34 ± 0.25 3.72 ± 0.27 3.46 ± 0.26 3.09 ± 0.15 3.02 ± 0.17 3.80 ± 0.16 3.50 ± 0.14 3.04 ± 0.15 3.27 ± 0.22 3.23 ± 0.22 3.13 ± 0.22 3.17 ± 0.16 3.54 ± 0.19 3.36 ± 0.18 3.50 ± 0.19 3.11 ± 0.17 2.98 ± 0.16 3.50 ± 0.18 3.49 ± 0.18 3.01 ± 0.22 3.18 ± 0.23 3.32 ± 0.24 3.23 ± 0.24 3.08 ± 0.17 3.00 ± 0.23 3.28 ± 0.18 3.46 ± 0.19 3.88 ± 0.20 3.78 ± 0.20 3.92 ± 0.26 3.78 ± 0.20 3.66 ± 0.20 3.57 ± 0.19 3.09 ± 0.15 3.14 ± 0.23 2.83 ± 0.20 3.38 ± 0.25 3.13 ± 0.23 3.00 ± 0.14 2.87 ± 0.34 2.86 ± 0.28

6a + 12b 6a + 12b 6a + 12b 6a + 12b 6a + 12b 6a + 12b 6a + 12b 6a + 12b 8a + 11b 6a + 12b 6a + 12b 6a + 12b 6a + 12b 6a + 12b 6a + 8b 6a + 12b 6a + 12b 6a + 14b 6a + 12b 6a + 12b 6a + 12b 6a + 12b 6a + 14b 6a + 14b 6a + 12b 6a + 12b 6a + 12b 6a + 12b 6a + 12b 6a + 12b 6a + 12b 6a + 12b 6a + 12b 6a + 12b 6a + 13b 6a + 14b 6a + 12b 6a + 13b 6a + 12b 6a + 12b 6a + 12b 6a + 12b 6a + 12b 6a + 12b 6a + 12b 6a + 12b 6a + 12b 6a + 12b 6a + 12b 6a + 12b 6a + 12b 6a + 14b 6a + 12b 6a + 12b 6a + 12b 6a + 12b 6a + 12b 6a + 14b 6a + 12b 6a + 12b 6a + 13b 6a + 12b 6a + 12b

0.078 ± 0.007 8.0 ± 0.3 9.7 ± 0.3 17.2 ± 0.4 18.4 ± 0.6 20.0 ± 0.8 23.3 ± 0.4 27.0 ± 0.7 40.3 ± 1.1 32.2 ± 0.8 27.8 ± 0.7 25.1 ± 0.7 40.4 ± 2.0 29.5 ± 1.1 42.6 ± 1.9 14.2 ± 0.4 9.2 ± 0.3 10.7 ± 0.2 13.7 ± 0.4 14.6 ± 0.4 19.1 ± 0.4 32.7 ± 1.0 24.6 ± 0.5 29.0 ± 0.7 28.2 ± 0.7 29.0 ± 0.5 31.6 ± 0.6 35.1 ± 0.9 31.4 ± 1.3 24.9 ± 0.3 26.8 ± 0.5 30.8 ± 0.6 32.9 ± 1.3 43.0 ± 1.3 6.3 ± 0.1 10.0 ± 0.25 16.7 ± 0.5 36.1 ± 0.9 54.1 ± 1.0 7.8 ± 0.2 16.9 ± 0.3 27.5 ± 0.7 26.6 ± 0.7 31.4 ± 0.6 36.6 ± 0.7 36.8 ± 1.3 36.3 ± 0.7 3.7 ± 0.1 6.6 ± 0.2 7.5 ± 0.2 9.9 ± 0.3 11.2 ± 0.7 12.9 ± 0.4 14.2 ± 0.3 23.9 ± 0.5 23.0 ± 0.4 22.6 ± 0.9 21.2 ± 0.4 17.9 ± 0.3 19.3 ± 0.4 25.7 ± 0.6 24.3 ± 0.5 22.9 ± 0.6

0.023 ± 0.002 2.8 ± 0.2 3.3 ± 0.2 5.0 ± 0.3 5.4 ± 0.3 6.9 ± 0.5 8.1 ± 0.5 9.6 ± 0.6 12.8 ± 1.1 11.4 ± 0.6 9.0 ± 0.5 9.0 ± 0.5 11.8 ± 1.1 10.2 ± 0.9 15.8 ± 1.0 4.2 ± 0.3 2.6 ± 0.2 3.3 ± 0.2 3.4 ± 0.3 4.0 ± 0.3 5.4 ± 0.4 10.1 ± 0.7 7.4 ± 0.6 7.8 ± 0.6 8.1 ± 0.6 9.4 ± 0.5 10.5 ± 0.6 9.2 ± 0.5 9.0 ± 0.5 8.2 ± 0.4 8.2 ± 0.6 9.5 ± 0.7 10.5 ± 0.8 13.6 ± 0.8 1.8 ± 0.1 3.0 ± 0.2 4.8 ± 0.3 11.6 ± 0.7 18.2 ± 1.0 2.2 ± 0.1 4.8 ± 0.3 9.1 ± 0.7 8.3 ± 0.6 9.5 ± 0.7 11.3 ± 0.9 12.0 ± 0.8 12.1 ± 0.9 1.1 ± 0.1 1.9 ± 0.1 1.9 ± 0.1 2.6 ± 0.2 2.9 ± 0.3 3.4 ± 0.2 3.9 ± 0.2 6.7 ± 0.4 7.4 ± 0.4 7.2 ± 0.6 7.5 ± 0.5 5.3 ± 0.4 6.2 ± 0.5 8.6 ± 0.5 8.5 ± 0.9 8.0 ± 0.8

The sample ID marked with * means measured with 90–125 μm fraction, and sample ID marked with ‘IRSL’ means measured by IRSL. a Aliquot number used for SAR. b Aliquot number used for SGC.

therefore the OSL dating is suitable for the aeolian sediments in the eastern QB. The ages of the underlying fluvial sand of XRH1 (9.4 ± 0.7 and 10.4 ± 0.6 ka), XRH2 (13.6 ± 0.8 ka) and XRH3 (18.2 ± 1.0 ka) demonstrate that these terraces of the Xiariha River might formed at LGM, deglaciation and early Holocene.

Most of the OSL ages are in stratigraphic order in a given section within errors, except six ages marked with underlines in Figure 4. In the XRH5 section, underestimation of the OSL ages of XRH5-10 and XRH5-11, where a few anthropogenic ash layers were found, might due to the partly reset of OSL signal by fire. In the lower part of the

Please cite this article as: Yu, L.P., Lai, Z.P., Holocene climate change inferred from stratigraphy and OSL chronology of aeolian sediments in the Qaidam Basin, northeastern Qinghai–Tibetan Plateau, Quaternary Research (2013), http://dx.doi.org/10.1016/j.yqres.2013.09.006

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Wintle, 2006), however, the depositional rate of paleosol in the eastern QB could be as high as 0.5–1.0 m/ka (e.g., the XRH1 section) resulting from the increased dust-retention ability from vegetation, and this can reduce the influence of pedogenesis. The influence of bioturbation on OSL dating of samples from this area could be minor, as the climate is arid and the pedoturbation process is slow, resulting in accessional paleosol. Aeolian activities in the eastern QB since the deglaciation

Figure 5. Comparison of De estimates by single aliquot regeneration (SAR) and by standard growth curve (SGC). The blue lines show ±10% error of the diagonal. (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)

SYK1 section, the ages of SYK1-E, SYK1-8 and SYK1-B are obviously inverted, and if they are excluded, the chronology of this section will be well fitted with the depth. In the XRH1 section, OSL age of XRH1-2 is treated as outlier and excluded. However, most of these ages are still in stratigraphic order in 2 σ errors. The penetration of the pedogenetic process can transform the underlying original deposition (e.g., the loess) into paleosol; however, it was deposited under the former arid climate. In this case, the OSL ages for the humid climate when paleosol developed was overestimated (Lai and Wintle, 2006). This is common on the Chinese Loess Plateau where depositional rate of the paleosol is as low as ~0.1 m/ka (Lai and

It is important to know the timing when the wide distributed aeolian sediments started to accumulation in the QB. According to Hao et al. (1998), the paleosols in the eastern QB started to accumulate during the last interglaciation at ca. 75–130 ka. However, our OSL ages from the XRH4 section at the same location show these aeolian sediments accumulated since ~12 ka, which is much younger. Yu and Lai (2012) proposed that aeolian sediments in the paleo-dunes overlying the loess in the eastern QB formed since ca. 12.4 ka, and was much younger than the TL ages of Zeng et al. (2003) which displayed the aeolian sands accumulated since the LGM in the same sites. The TL ages in Hao et al. (1998) were much older than the OSL ages in XRH4 section from the same site, suggesting possible overestimation of TL ages (Yu and Lai, 2012) again. The accumulation of aeolian sediments started at ~12.1 ka according to the dating results in this study, which supports the conclusion that the aeolian sediments in the eastern QB were formed since the deglaciation (Yu and Lai, 2012; Yu et al., 2013). However, aeolian sediments accumulated during and before the LGM have been dated in many regions in the QTP, e.g., the Lhasa River region (Lai et al., 2009), the Qinghai Lake Basin (Liu et al., 2012), and the Gonghe Basin (Qiang et al., 2013). According to Sun et al. (2007), present loess deposits in the interior of Tibet accumulated after the last deglaciation and had a basal age of 13–11 ka due to erosion by glaciofluvial outwash during the beginning of the interglaciation period. OSL chronology in this study (13.6 ±

Figure 6. OSL growth curves, decay curves for weakly expressed soil sample XRH4-2 and aeolian sand sample XRH4-8.

Please cite this article as: Yu, L.P., Lai, Z.P., Holocene climate change inferred from stratigraphy and OSL chronology of aeolian sediments in the Qaidam Basin, northeastern Qinghai–Tibetan Plateau, Quaternary Research (2013), http://dx.doi.org/10.1016/j.yqres.2013.09.006

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L.P. Yu, Z.P. Lai / Quaternary Research xxx (2013) xxx–xxx

0.8 ka in the XRH2 section and 18.2 ± 1.0 ka in the XRH3 section) offer evidence for the fluvial processes, which can cause erosion to the aeolian sediments in the QB during LGM and deglaciation. In this study, the depositional rates of the paleosols are even much higher than that of the loess in some certain sections, e.g., XRH1 and XRH5, which is different from that of the Loess Plateau, where higher depositional rate is the characteristic of the loess (Lai and Wintle, 2006). This may denote that vegetation could be crucial for the aeolian sediment accumulation in the arid regions to trap the dust. This may explain the fact that the aeolian sediments were limited before LGM due to the poor vegetation. The earliest aeolian sand accumulated at 12.1 ± 0.7 ka in XRH4 section and at 11.4 ± 0.5 ka in XRH1 section, which might correspond to the YD event. The weakly expressed soil and paleosols occurred at ~11.6 ka (at 11.8 ± 1.1 ka in SYK1, before 10.4 ± 0.8 ka in XRH2, at 11.6 ± 0.7 ka in XRH3 and at 11.3 ± 0.6 ka in XRH4), which reflect the increase of the effective moisture at the beginning of the Holocene. The development of paleosols ended mainly before the early Holocene. The aeolian sand and loess deposit started asynchronous in different sites during the mid- to late Holocene, and lasted till the modern time. In general, paleosols mainly developed during the early and midHolocene, while the loess and aeolian sand mainly accumulated at midto late Holocene, demonstrating a drying trend during the Holocene. Discussion

e.g., 3.3 ± 0.2 ka and 3.4 ± 0.3 ka in XRH1 section, and the boundary age is easy to identify. If the errors are larger, e.g., 9.2 ± 0.5 ka and 9.6 ± 0.6 ka in the WLS1 section, the boundary age is estimated as 9.4 ka with only half a std deviation used. Aeolian sediments are easy to be eroded, and are therefore usually discontinuous, however, with our high resolution OSL chronology, the hiatuses are easily to be identified. To ensure the reliability of the EMI, the periods of hiatuses are not evaluated. After the boundary ages are calculated, an individual sedimentary sequence can be displayed as an EMI curve, though it might not be continuous due to the existence of hiatuses. Then these EMI curves could be averaged to obtain a uniform mean EMI curve (Fig. 7A), which presents general effective moisture change of this region. Because the boundary ages are of centurial resolution, the final mean EMI curve is also with a 100-year interval. Except for the sections in this study, another six former reported loess–aeolian sand sections (WLS1, BDB1, BDB1-2, XXT2 and XXT2-2 in Yu and Lai (2012), and HBC1 section (unpublished data), Figs. 1C and 4) are also included. Taking into account the basal ages and their errors of these aeolian sections, the EMI was reconstructed since 12.8 ka (Fig. 7A), and this is also the time when the YD event started. EMI change since 12.8 ka in the eastern QB The stratigraphy, in general, demonstrates a drying trend during the Holocene. This drying trend is more obviously illustrated by the EMI curve in Figure 7A. The EMI change since 12.8 ka can be separated into four periods: 12.8–11.6 ka, 11.6–8.3 ka, 8.3–3.5 ka and 3.5–0 ka.

Effective moisture change since 12.8 ka in the eastern QB Effective moisture index The purpose of this study is to reconstruct paleoclimatic change, especially to identify source of the moisture, in the eastern QB during the Holocene. Accordingly, a uniform paleoclimatic record of the eastern QB is needed to have a better comparison with records from the surrounding regions. However, a series of aeolian sections with many OSL ages and stratigraphic units (Fig. 4) is difficult to make a visualized comparison. In this study, we try to apply the effective moisture index (EMI) method to the aeolian sediments, which has been successfully applied to lacustrine sediments to get a uniform paleoclimatic record from a certain region and make relevant comparison (e.g., Chen et al., 2008; Zhao et al., 2009a,b). The basic promise of EMI for aeolian sediments is that the evolution of aeolian sediments is mainly controlled by the effective moisture change. Although aeolian sand activity is related to transport capacity of the wind (Lancaster, 1988) and sand supply (Cohen et al., 2010), generally speaking, the formation of aeolian sand is a response to arid climate, while paleosol developed under a relatively humid climate after the dunes were stabilized in north China (e.g., Zhu et al., 1980; Liu, 1985; Sun et al., 1996; Yang et al., 2011). In the Qinghai Lake region, northeastern QTP, alternate deposition of aeolian sediments were also controlled by the effective moisture (Lu et al., 2011). In the hyper-arid QB, there is enough sand supply from the yardangs, playas and alluvial fans, and the wind transport ability is constantly strong during the winter and spring. Therefore, effective moisture change, which can influence the vegetation and pedogenesis, might be the crucial factor for the aeolian sediments evolution. Based on the aforementioned hypothesis, the EMI is applied to aeolian sediments to demonstrate the average condition of the selected profiles during the Holocene. The aeolian sand, loess, weakly expressed soil and paleosols in the stratigraphy receive EMI scores of 1, 2, 3 and 4, respectively. EMI = 1 stands for the lowest effective moisture, and 4 stands for the highest. The most important process during the EMI construction is to calculate the boundary ages between different stratigraphic units, which are based on direct OSL dating or indirect depositional rate. The errors of most OSL ages in this study vary between 0.1 and 0.9 ka, therefore, the boundary ages are in centurial scale. In most cases, OSL ages over and beneath a certain boundary are very close,

12.8–11.6 ka (Deglaciation): hyper-arid. EMI record is the lowest during 12.8–11.6 ka, which is calculated from the aeolian sand in SYK1 section (11.4 ± 0.5 and 11.5 ± 0.5 ka), XRH4 section (12.0 ± 0.7 and 12.1 ± 0.7 ka), XXT2 section (12.4 ± 0.7 ka), XXT2-2 section (11.6 ± 0.9 ka), BDB1 section (11.7 ± 0.6 ka), and BDB1-2 section (11.5 ± 0.6 ka). This hyper-arid period of 12.8–11.6 ka might correspond to the worldwide YD event, though several ages are around 11.6 ka (11.4 ± 0.5, 11.5 ± 0.5, 11.6 ± 0.9, 11.7 ± 0.6, and 11.5 ± 0.6 ka) at the end of the YD event. That is because most of the OSL samples were taken near the upper boundaries to display when these units terminate, and these ages obviously show this arid period terminated at ca. 11.6 ka. The YD event was found at 12.70–11.55 ka in GRIP ice core (Johnsen et al., 1992) and at 12.89–11.65 ka in GISP2 ice core (Alley et al., 1993) in Greenland, and was also detected in Chaka salt lake sediments in the eastern QB (Liu et al., 2008), in Qinghai Lake (7 in Fig. 1B) sediments on the northeastern QTP (An et al., 2012b) and in the Guliya ice core on the QTP (Thompson et al., 1997). The other possibility is that this arid period is not corresponding to the YD event alone, but to a later part of the consistent arid stage from the last glacial to late deglaciation, if the ASM was not strong enough to penetrate to the QB during the deglaciation. From a core in the Qinghai Lake, east to the QB, An et al. (2012b) inferred that the summer monsoon index (SMI) is relatively low and has lower-amplitude fluctuations during 32–11.5 ka, because the monsoon front rarely penetrated sufficiently northwest to reach the Qinghai Lake region, at the fringes of the modern ASM. The eastern QB located at the fringes of ASM as well, and the effective moisture might be consistently low before 11.6 ka, and therefore it is difficult to infer whether this arid period is the response to the YD event alone until paleosol or loess presenting higher EMI is found before 12.8 ka. During this period, aeolian sand accumulated in many sections, and there is no evidence of loess and paleosols developed, demonstrating an arid environment with poor vegetation. Aeolian sediments formed during this period are mainly sand and loess, without paleosols in the Qinghai Lake region (Lu et al., 2011; Liu et al., 2012). 11.6–8.3 ka (Early Holocene): humid and variable. The effective moisture record has an extremely abrupt transition at ~11.6 ka at the beginning of the Holocene, like the SMI record in the Qinghai Lake region (An et al.,

Please cite this article as: Yu, L.P., Lai, Z.P., Holocene climate change inferred from stratigraphy and OSL chronology of aeolian sediments in the Qaidam Basin, northeastern Qinghai–Tibetan Plateau, Quaternary Research (2013), http://dx.doi.org/10.1016/j.yqres.2013.09.006

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Figure 7. (A) Effective moisture index (EMI) of the eastern Qaidam Basin (QB) since the Younger Dryas (YD) event and (B) comparison with insolation change of the 36°N, the latitude of the QB (Berger and Loutre, 1991), (C and D) δ18O records in Dongge Cave (Dykoski et al., 2005; Wang et al., 2005), (E) average moisture index for arid central Asia (Chen et al., 2008), (F) δ13C records from Hongyuan peat in the eastern QTP (Hong et al., 2003), (G) Summer Monsoon Index (SMI) of the Qinghai Lake sediments (An et al., 2012b) and (H) δ18O records in Qunf Cave, Oman (Fleitmann et al., 2003).

2012b). The effective moisture reached the highest level during the early Holocene, and the fluctuations were also larger. There is growing evidence for a strong ASM on the QTP in the early Holocene (e.g., Lister et al., 1991; Hong et al., 2003; Shen et al., 2005; Zhao et al., 2009a; An et al., 2012b). The humid and warm early Holocene was beneficial to the development of the paleosols, with both the strengthened pedogenesis and increased dust-retention ability resulting from the development of vegetation. In the Qinghai Lake region, paleosols also occurred during the early Holocene with the increase effective moisture (Lu et al., 2011; Liu et al., 2012). 8.3–3.5ka (Mid-Holocene): relatively humid and stable. The humid period terminated with the coming of the global 8.2 ka cooling event, which displayed as a drying period with low EMI during 8.3–8.1 ka in this study. This drying period is displayed apparently by the intercalation of aeolian sand into the loess/paleosols sediments in two sites of this study, e.g., 8.1 ± 0.5 ka in SYK1 section and 8.2 ± 0.4 ka in XRH2 section. This worldwide Holocene cooling event, possibly related to abrupt

outflow from a Laurentide ice-marginal lake (Barber et al., 1999), was represented in deep-sea cores (Bond et al., 2001), Greenland ice cores (Dansgaard et al., 1993; Grootes et al., 1993) and many other kinds of paleoenvironmental archives (e.g., Neff et al., 2001; Hong et al., 2003; Wang et al., 2005; Mischke and Zhang, 2010; Hou and Fang, 2011). The effective moisture decreased dramatically at 8.3–8.1 ka and increased after that epoch. The effective moisture during 8.3–3.5 ka was generally stable with minor fluctuation, but it never reached the level of the early Holocene. This moderately humid and stable condition was suitable for the development of paleosols and loess in the eastern QB. The loess overlying paleo-dunes are mainly formed during the mid-Holocene (Yu and Lai, 2012). Paleosols in the Qinghai Lake region also mainly developed during the mid-Holocene (Liu et al., 2012). 3.5–0 ka (Late Holocene): drying. The effective moisture decrease in late Holocene started at 3.5 ka. In the middle QB, lake level of the Xiao Qaidam Lake decreased ca. 10 m since 3 ka (Sun et al., 2010), and the lake level of the Qinghai Lake dropped more than 8 m after 3.7 ±

Please cite this article as: Yu, L.P., Lai, Z.P., Holocene climate change inferred from stratigraphy and OSL chronology of aeolian sediments in the Qaidam Basin, northeastern Qinghai–Tibetan Plateau, Quaternary Research (2013), http://dx.doi.org/10.1016/j.yqres.2013.09.006

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0.4 ka (Liu et al., 2011) with the increase of the aridity. Strongly enhanced aridity at this time is also indicated by the δ18O records in Dongge Cave (Dykoski et al., 2005, Fig. 7C) and Qunf Cave in Oman (Fleitmann et al., 2003, Fig. 7H), ISM records in the Qinghai Lake (An et al., 2012b, Fig. 7G), and peat records in Hongyuan (Hong et al., 2003, Fig. 7F). Under this arid climate, aeolian activities increased in the eastern QB, e.g., sediments were mainly aeolian sand and loess during the late Holocene, instead of paleosols during the early and mid-Holocene. Moreover, the paleo-dunes in the middle and southwestern margin accumulated at ca. 3–4 ka (Yu et al., 2013), aeolian sand layers formed at ca. 3.4– 3.2 ka in the southern margin (Zeng, 2006), and linear dunes developed at 3.2–0.8 ka in the central QB (Zhou et al., 2012) were all related to the arid climate in the late Holocene. In addition, the climate deterioration around 4 ka led to the collapse of Neolithic culture in the western part of the Chinese Loess Plateau, which is close to the eastern part of the QB (An et al., 2005). Moisture source of the eastern QB The total organic carbon, total nitrogen and mineralogy records from a core at Chaka salt lake, located to the east of the QB, illustrated that the period of 11.4–5.3 ka was warm and humid but that after 5.3 ka it was cold and dry (Liu et al., 2008). The study from the Da Qaidam and Bieletan saline lakes in the middle QB revealed that the salinity in the lakes decreased starting at ~11.5 cal ka BP, and the effective humidity peak occurred at 11.5–8.9 cal. ka BP with the increased temperature and precipitation (Huang et al., 1981). Sun et al. (2010) indicated that the lake level of the Xiao Qaidam Lake (2 in Fig. 1B) was ~10 m higher than the present at ca. 11–3 ka and dropped after that time. Similar to our results, these three lacustrine records showed increase of precipitation at the beginning of the Holocene (~11.5 ka) and a humid early to mid-Holocene from the eastern to middle QB. However, fossil pollen data, oxygen isotope and Mg/Ca ratios from Hurleg Lake in the eastern QB indicate a dry climate in the early and mid-Holocene, followed by a wet climate in the mid- to late Holocene there (Zhao et al., 2007). There is growing evidence for a strong Asian summer monsoon on the QTP in the early Holocene, associated with a peak of Northern Hemisphere summer insolation. The SMI record (An et al., 2012b; Fig. 7G), the pollen diagram (Shen et al., 2005), the δ18O records of ostracode shells (Lister et al., 1991; Lui et al., 2007; An et al., 2012b), redness index (Ji et al., 2005) from lacustrine sediments, and the aeolian sediments from Halali section (Chen et al., 1991) all reflected a humid early to mid-Holocene and a arid late Holocene in the Qinghai Lake area. Moreover, Holocene effective moisture change on the QTP synthesized by Zhao et al. (2009a) on the basis of six records (see the references therein) suggested similar trend. The climate during the early Holocene was humid in regions influenced by the ASM (Fleitmann et al., 2003, Fig. 7H; Dykoski et al., 2005, Fig. 7C; Wang et al., 2005, Fig. 7D; Shao et al., 2006; Hu et al., 2008), while it was arid in regions controlled by the westerlies (Chen et al., 2008, Fig. 7E). The similarity of our results (Fig. 7A) to that of the regions affected by the ASM (e.g., Dykoski et al., 2005; Wang et al., 2005; Figs. 7 C and D) demonstrates that the Holocene moisture change in the eastern QB was controlled by the ASM. However, it is difficult to identify whether the precipitation in the eastern QB was mainly controlled by the EASM or the ISM. Forcing mechanism of effective moisture change in eastern QB during the Holocene The Holocene effective moisture change in the northeastern QTP was controlled by the intensity of the ASM, and the ASM change was mainly controlled by the insolation during both the Holocene (Wang et al., 2005) and deglaciation (An et al., 2012b). Consequently, the EMI change in the eastern QB was mainly controlled by the Northern

Hemispheric summer insolation, and our results indicate that the effective moisture change in the eastern QB was parallel to the insolation change of 30°N (Fig. 7B). During 12.8–11.6 ka at the late deglaciation, the ASM might not be strong enough to reach the QB, and with the influence of the global cooling YD event, the eastern QB was hyper-arid. At the beginning of the Holocene, the insolation was very high, and the ASM strengthened and penetrated into the QB which increased the precipitation abruptly. The abrupt increase of both EMI in the eastern QB in this study and SMI in the Qinghai Lake area (An et al., 2012b; Fig. 7G) was mainly because these regions were beyond the influence of ASM during the deglaciation, and a sudden occurrence of intensive ASM might result in an abrupt change in the paleoclimatic records. The Dongge Cave was influenced by the ASM consistently, and its response to the increase of ASM was more gradual (Dykoski et al., 2005; Fig. 7C). Strong summer insolation in the Northern Hemisphere during the early Holocene (Berger and Loutre, 1991) induced strong land–ocean pressure and temperature gradients, and this increased onshore flow of moist air in summer, causing an enhanced ASM (COHMAP, 1988). The gradual weakening of the ASM since the mid-Holocene was in response to the decrease in summer insolation (Gupta et al., 2003), and enhanced by the feedbacks from change in vegetation cover and soil moisture (Kutzbach et al., 1996). Although the decrease of distance to the oceanic moisture source had increased the intensity of the ASM (An et al., 2012b), variability of the sea level was also controlled by the insolation change. Lacustrine records from the Qinghai Lake revealed that the warmhumid ASM climates were anti-correlated with cold–dry westerlies climate, and that the alternating influence of the westerlies and the ASM on glacial–interglacial time scales likely had been a major pattern of climate change in these regions throughout most of the Quaternary (An et al., 2012b). This reveals the waxing and waning relationship between the ASM and the westerlies in the northeastern QTP. In the eastern QB when the ASM shrank, the cold–dry westerlies might have strengthened, increasing the arid trend during the late Holocene. The Holocene EMI change in the eastern QB was controlled by the strength of ASM, which generally follows the orbitally induced Northern Hemispheric insolation change. With the weakening of the ASM, the westerlies strengthened and promoted the aeolian activities in the eastern QB. Conclusions (1) The surface aeolian sediments accumulated at least before 12.4±0.7ka during the deglaciation. The paleosols mainly developed during early and mid-Holocene, while the aeolian sand and loess mainly accumulated during the mid- and late Holocene. The underlying fluvial sediments mainly deposited at the LGM, deglaciation and early Holocene. (2) The method of EMI based on stratigraphy and high-resolution chronology is suitable for aeolian sediments to demonstrate overall paleoclimatic information in the arid zone. The effective moisture change revealed by the EMI in the eastern QB during the Holocene could be summarized as: a humid and fluctuant early Holocene at 11.6–8.3 ka, a moderate humid and stable mid-Holocene at 8.3–3.5ka, and an arid late Holocene at 3.5–0ka. (3) The effective moisture change in the eastern QB was mainly controlled by intensity of the ASM, which generally followed the orbitally induced Northern Hemispheric summer insolation change. Acknowledgments This study was supported by China NSF (41290252), SKLLQG (SKLLQG1217), Qinghai Science & Technology Department (2013-Z-943Q), and China Postdoctoral Science Foundation project

Please cite this article as: Yu, L.P., Lai, Z.P., Holocene climate change inferred from stratigraphy and OSL chronology of aeolian sediments in the Qaidam Basin, northeastern Qinghai–Tibetan Plateau, Quaternary Research (2013), http://dx.doi.org/10.1016/j.yqres.2013.09.006

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Please cite this article as: Yu, L.P., Lai, Z.P., Holocene climate change inferred from stratigraphy and OSL chronology of aeolian sediments in the Qaidam Basin, northeastern Qinghai–Tibetan Plateau, Quaternary Research (2013), http://dx.doi.org/10.1016/j.yqres.2013.09.006